SUPPLEMENTARY INFORMATION

Transcription

1 SUPPLEMENTARY INFORMATION Extreme deepening of the Atlantic overturning circulation during deglaciation Stephen Barker, Gregor Knorr, Maryline Vautravers, Paula Diz and Luke Skinner Supplementary Information S1. Modern deep sea ventilation at TNO57-21 Figure S1 shows GLODAP 1 data for a meridional section through the Atlantic. The site of TNO57-21 lies in the mixing zone between well-ventilated deep waters originating in the north and poorly-ventilated deep waters that surround Antarctica. Deep water chemistry at the site is currently dominated by the southerly end member leading to relatively low values of [CO = 3 ], Δ 14 C and δ 13 C. S2. Deep water ventilation from benthic foraminiferal 14 C Benthic 14 C ages may be employed to reconstruct deep water ventilation age by two alternative approaches. The benthic-planktonic (B-P) age offset describes the contemporary 14 C gradient between surface and bottom waters. However, as pointed out by Adkins et al. 2, changes in atmospheric Δ 14 C mean that the instantaneous offset may be misleading. The projection age method 2 takes this into account by combining a benthic 14 C age with its known calendar age and back-projecting along the 14 C decay curve until it intersects with the known record of atmospheric Δ 14 C (or in this case the tropical surface ocean 3 ) (Fig. S2). The projection age method requires knowledge of the absolute (calendar) age of the sample and the record of atmospheric Δ 14 C (or equivalent). Calendar ages for our samples were reported previously 4. Here we use the record of Δ 14 C determined for the sub-tropical ocean as reported by Hughen et al. 3,5 to calculate the projection age (Table S1). This is equivalent to using the atmospheric record (which is anyway derived from that used here over the interval of interest) and allows better comparison with the B-P (benthic-planktonic) ages. The modern projection age at the site of TNO57-21 is 200 yr greater than the equivalent B- nature geoscience 1

2 supplementary information P age (main text Fig. 1). This reflects the 200 yr reservoir correction applied to planktonic 14 C ages at this site 4. S3. Carbonate dissolution in TNO57-21 S3.1. Reconstructing carbonate dissolution There are several ways of reconstructing the history of carbonate dissolution at the seafloor although each has its particular complexities 6. In order to assess changes in the state of preservation in a marine core it is sensible to compare various proxies in order to assess the possible influences of e.g. changing primary environmental conditions (such as production at the sea surface) versus the effects of dissolution and dilution at the seafloor. We have measured a number of common proxies for carbonate dissolution in TNO57-21, including %coarse fraction (%>63μm), foraminiferal fragmentation and whole foraminiferal shells per gram (the latter two being counted in the >150μm fraction) and compare these with the record of %CaCO 3 from the same core, as reported by Sachs and Anderson 7 (Fig. S3). The records reveal commonalities and differences that reflect the various auxiliary controls and influences on each proxy. All proxies suggest better preservation during the B-A and D-O 8. But while the records of fragmentation, coarse fraction and intact shells per gram display prominent peaks at these times, the record of %CaCO 3 reveals two broad maxima. These appear not to be driven by increased surface ocean primary productivity as reconstructed by Sachs and Anderson 8 in the same core (see also below). Given the rather low coarse fraction content of this core (typically 2-3 % >63μm) the CaCO 3 content must be driven primarily by changes within the fine fraction (<63μm), which itself can be influenced by lateral transport (the site of TNO57-21 is a drift deposit). As discussed by Sachs and Anderson 7 the alkenone signal within TNO57-21 (which is also derived from the fine fraction) involves transport of the host material from several degrees of latitude to the north. Such transport could then have smeared out some of the transient changes we observe in our other dissolution proxies. The records of %CaCO 3, coarse fraction and whole shells per gram all suggest very poor preservation during the interval ~32-22 ka ago but although the record of fragmentation 2 nature geoscience

3 supplementary information generally agrees with our other measured proxies it does not show particularly high values throughout this interval. We counted fragments in the same size fraction as the faunal assemblage counts (as described by 4 ) (i.e. >150μm). For a large extent of dissolution we would expect continued break-up of fragments into smaller and smaller size fractions. The very low absolute numbers of fragments between ~32-22 ka ago (Fig. S3) may reflect the loss of fragments into the fraction <150μm through continued break up. It is also possible that changes in assemblage have a greater influence on fragmentation (perhaps by the number of fragments any particular shell breaks into) when dissolution is extreme. In any case it is possible that the record of fragmentation is not giving a truly representative record of dissolution for the period concerned. Again this highlights the value of using multiple proxies for reconstructing dissolution. We note that our use of fragmentation for quantifying dissolution in our recent study of surface records from TNO is unlikely to be problematic because the key interval of importance for that study was HS1, when fragmentation agrees well with both %coarse fraction and whole shells per gram. Changes in productivity could influence our records in two main ways. Firstly an increase in productivity could lead to more carbonate production and hence higher %CaCO 3 in the core. However, from Figure S3 it is apparent that productivity, as reconstructed using the concentration of authigenic U 8, is sometimes high when %CaCO 3 is low while at other times it is low when %CaCO 3 is low. Alternatively increased productivity could drive increased dissolution by supplying more organic carbon to the sediments and lowering ambient [CO = 3 ] within the pore waters 9. Since this is a very deep site we may expect bottom water [CO = 3 ] (rather than the flux of organic carbon) to dominate the degree of dissolution within TNO Furthermore, production was not particularly low during the B-A as compared with the Holocene for example. Therefore we conclude that dissolution at our site is driven predominantly by deep water changes rather than surface ocean productivity. One then might ask why there is any correlation at all between surface ocean production and carbonate dissolution at our site. We would argue that it is precisely the location of TNO57-21 (at the northern margin of the Southern Ocean) that results in its nature geoscience 3

4 supplementary information sensitivity to essentially simultaneous changes in the both the surface and deep ocean as a result of changes in the AMOC and the bipolar seesaw. S3.2. Timescales of changing dissolution Berger 11 described a deglacial preservation spike in both the Atlantic and Pacific Oceans. Although the original paper does not show the data in detail the author describes the event as a relatively short (~1000 yr) event occurring during the maximum rate of change in the oxygen isotope signal. This would place the event during the deglaciation although we cannot be sure that it is the same event as we describe here. Broecker et al. 12 also describe a transient carbonate preservation event that might be expected due to the re-growth of boreal forests after the last ice age. In this scenario the preservation event would be a broad maximum occurring during the early Holocene; not the transient feature we observe during deglaciation. The early Holocene preservation maximum described by Broecker et al. 12 would be expected to lead to a subsequent decrease in whole ocean [CO = 3 ] due to carbonate compensation (the tendency toward steady state between input and output of alkalinity). This would occur over a period of several thousands of years Indeed a widespread increase in carbonate dissolution has been observed during the last yr 17 and our records also suggest worsening carbonate preservation throughout the course of the Holocene (Fig. S3). However, both the onset and decline of the sharp preservation events we observe during the B-A and D-O 8 occur on timescales of hundreds of years (not thousands). Furthermore, it seems most likely that both of the preservation events we observe were caused by a similar mechanism, which presumably rules out processes which were specific to deglaciation. Thus we conclude that these events were not a result of whole ocean chemical variations but must have been driven by alternation between water masses with very different chemistries. S4. Ventilation of the Southern Ocean during Heinrich Stadial events Two recent studies have argued for the release of CO 2 from within the Southern Ocean during Heinrich Stadial events 4,18. Ventilation of the Southern Ocean during these events should drive an increase in its bulk [CO = 3 ] by the release of CO 2 but we see strong 4 nature geoscience

5 supplementary information dissolution at our site at least during the latter part of HS1. This is not contradictory to expectations. The release of old, corrosive waters from within the Southern Ocean during HS1 would of course cause an improvement of ventilation of this water mass but would also lead to worsened ventilation in areas sourced from the region of release. Particularly 14 C-depleted intermediate waters observed off Baja, California during HS1 19 may indeed reflect the release of poorly ventilated waters at this time. Since our site is on the margin of the Southern Ocean we suggest that an increase in dissolution is a natural consequence of Southern Ocean ventilation and release of corrosive waters during HS1 (and presumably during other HS events). S5. Model simulation of an AMOC overshoot at the B/A transition S5.1. Model Set-up The three-dimensional ocean general circulation model is based on the large-scale geostrophic model (LSG) 20. The horizontal resolution is 3.5 on a semi-staggered grid (type E ) with 11 levels in the vertical. The model includes a simple thermodynamic sea ice model, a 3rd order advection scheme for temperature and salinity 21, and a parameterization of overflow 22,23. For glacial conditions the storage of water in inland ice sheets is taken into account by setting all ocean points with present-day water depth less than 120 m to land. This causes the Bering Strait to be closed and shallow areas like the Arctic shelves to become land. The ocean is driven by monthly fields of wind stress, surface air temperature and freshwater flux, which are taken from a present-day and a LGM simulation of the atmospheric general circulation model ECHAM3/T42 24,25. In order to close the hydrological cycle, a runoff scheme transports freshwater from the continents to the ocean. We employ a modeling approach, which allows an adjustment of surface temperatures and salinity to changes in ocean circulation, based on an atmospheric energy balance model. In the model, sea surface salinity can freely evolve. The heat flux exchange Q at the ocean surface is formulated as suggested by 26 : Q = ( λ 1 λ 2 2 ) (T a -T s ) (1) nature geoscience 5

6 supplementary information where T a is the prescribed air temperature, and T s denotes the ocean surface temperature. The heat-transfer coefficients λ 1 and λ 2 are chosen to be 15 W m 2 K 1 and W K 1, respectively. The applied heat flux parameterization has shown to be a suitable choice, allowing the simulation of observed SSTs and the maintenance of large-scale temperature anomalies in perturbation experiments 27. For the present-day state, we obtain a THC with a export of 14 Sv (1 Sv = 10 6 m 3 s 1 ) at 30 S and maximum heat transport of 1.1 PW, which is in the range of observations 28. For the LGM, our modelled ocean circulation is characterized by a weaker circulation, which is in agreement with proxy data [e.g., 29 ]. The glacial overturning circulation in the Atlantic can be characterized by a export of 8.5 Sv and southward shifted convection sites compared to present-day sites 30,31. is formed in the subpolar North Atlantic and the North Atlantic Current flows in a zonal direction along the horizontal density gradients, consistent with reconstructions [e.g., 32 ]. A detailed description of the modern THC and hydrographic fields can be found elsewhere 27,33. S5.2. Design of the transient deglaciation scenario The setup refers to experiment B4 of 34 and our Figure 4 shows the modelled deep Atlantic flow fields after 6000 (Fig. 4a) and 6200 (Fig. 4b) model years respectively. The gradual global deglaciation is implemented in the model by a linear transition from global glacial to interglacial background climate conditions in temperature, sea ice in the Southern Ocean and wind stress. The respective changes at each grid point are prescribed according to a linear trend between its glacial and its interglacial value, which gives rise to a linear progression between two temperature, sea ice and wind stress distributions. The freshwater fluxes remain constant in glacial conditions in order to isolate the deglacial temperature and meltwater impact on the AMOC. For a more detailed description of the forcing fields we would like to refer to 25. The model forcing at a certain time is given by CLIMATE FORCING = LGM + (PD LGM) t/t (2) 6 nature geoscience

7 supplementary information where LGM and PD are glacial and pre-industrial interglacial climate background conditions, t the integration time in model years and T the time of a full glacial to interglacial transition. We haven chosen T = 15,000 a, which has been motivated by longterm temperature trends as recorded in ice cores from Greenland and Antarctica during termination I [e.g. 35. In the Northern Hemisphere the ocean model includes a simple thermodynamic sea ice model. In our deglacial scenario the climate forcing is held constant after 7000 model years in order to obtain an equilibrium climate state. On the basis of geological evidence potential meltwater pulses have been suggested to occur at 19 ka B.P. (19 ka MWP) 36, 17.5 ka B.P. (H1b) and 16 ka B.P. (H1a) 37. The 19 ka MWP is simulated by a freshwater-flux magnitude of 0.25 Sv 36 with a duration of 200 a. Both Heinrich pulses H1a and H1b are performed with a 400 a freshwater release of 0.15 Sv centered at 17.5 and 16 ka. A deglacial background meltwater discharge of 0.05 Sv is applied throughout the remaining time of experiment between 1000 and 8000 model years. All deglacial meltwater discharge to the North Atlantic is uniformly distributed between 20 N and 50 N. Supplementary references 1. Key, R. M. et al. A global ocean carbon climatology: Results from Global Data Analysis Project (GLODAP). Glob. Biogeochem. Cycle 18, GB4031 (2004). 2. Adkins, J. F. & Boyle, E. A. Changing atmospheric ΔC 14 and the record of deep water paleoventilation ages. Paleoceanography 12, (1997). 3. Hughen, K. A. et al. Marine04 marine radiocarbon age calibration, 0-26 cal kyr BP. Radiocarbon 46, (2004). 4. Barker, S. et al. Interhemispheric Atlantic seesaw response during the last deglaciation. Nature 457, (2009). 5. Hughen, K. et al. 14 C activity and global carbon cycle changes over the past 50,000 years. Science 303, (2004). 6. Barker, S. in Encyclopedia of Quaternary Science (ed. Elias, S. A.) 1711 (Elsevier, 2006). 7. Sachs, J. P. & Anderson, R. F. Fidelity of alkenone paleotemperatures in southern Cape Basin sediment drifts. Paleoceanography 18, art. no (2003). 8. Sachs, J. P. & Anderson, R. F. Increased productivity in the subantarctic ocean during Heinrich events. Nature 434, (2005). nature geoscience 7

8 supplementary information 9. Emerson, S. & Bender, M. Carbon fluxes at the sediment-water interface of the deep sea: Calcium carbonate preservation. J. Mar. Res. 39, (1981). 10. Archer, D. Modeling the calcite lysocline. J. Geophys. Res.-Oceans 96, (1991). 11. Berger, W. H. Deep-Sea Carbonate and Deglaciation Preservation Spike in Pteropods and Foraminifera. Nature 269, (1977). 12. Broecker, W. S. et al. A search for an Early Holocene CaCO 3 preservation event. Paleoceanography 8, (1993). 13. Broecker, W. S. & Peng, T. H. The role of CaCO 3 compensation in the glacial to interglacial atmospheric CO 2 change. Glob. Biogeochem. Cycle 1, (1987). 14. Keir, R. S. On the late Pleistocene ocean geochemistry and circulation. Paleoceanography Paleoceanography 3, (1988). 15. Sundquist, E. T. Influence of Deep-Sea Benthic Processes on Atmospheric Co2. Philos. Trans. R. Soc. Lond. Ser. A-Math. Phys. Eng. Sci. 331, (1990). 16. Archer, D., Kheshgi, H. & Maier-Reimer, E. Multiple timescales for neutralization of fossil fuel CO2. Geophys. Res. Lett. 24, (1997). 17. Broecker, W. S. & Clark, E. Holocene atmospheric CO 2 increase as viewed from the seafloor. Glob. Biogeochem. Cycle 17, art. no (2003). 18. Anderson, R. F. et al. Wind-Driven Upwelling in the Southern Ocean and the Deglacial Rise in Atmospheric CO 2. Science 323, (2009). 19. Marchitto, T. M., Lehman, S. J., Ortiz, J. D., Fluckiger, J. & van Geen, A. Marine radiocarbon evidence for the mechanism of deglacial atmospheric CO 2 rise. Science 316, (2007). 20. Maier-Reimer, E., Mikolajewicz, U. & Hasselmann, K. Mean circulation of the Hamburg LSG OGCM and its sensitivity to the thermohaline surface forcing. J. Phys. Oceanogr. 23, (1993). 21. Schäfer-Neth, C. & Paul, A. in The Northern North Atlantic: A Changing Environment (eds. Schäfer, P., Ritzrau, W., Schluter, M. & Thiede, J.) (Springer, Berlin, 2001). 22. Lohmann, G. The influence of a near-bottom transport parameterization on the sensitivity of the thermohaline circulation. J. Phys. Oceanogr. 28, (1998). 23. Lohmann, G. & Schulz, M. Reconciling Bolling warmth with peak deglacial meltwater discharge. Paleoceanography 15, (2000). 24. Roeckner, E. et al. Simulation of the present-day climate with the ECHAM model: Impact of model physics and resolution. MPI Report No. 93 (Max-Planck-Institut für Meteorologie, Hamburg, 1992). 25. Lohmann, G. & Lorenz, S. On the hydrological cycle under paleoclimatic conditions as derived from AGCM simulations. J. Geophys. Res.-Atmos. 105, (2000). 26. Rahmstorf, S. & Willebrand, J. The Role of Temperature Feedback in Stabilizing the Thermohaline Circulation. J. Phys. Oceanogr. 25, (1995). 27. Prange, M., Lohmann, G. & Paul, A. Influence of vertical mixing on the thermohaline hysteresis: Analyses of an OGCM. J. Phys. Oceanogr. 33, (2003). 8 nature geoscience

9 supplementary information 28. Macdonald, A. M. & Wunsch, C. An estimate of global ocean circulation and heat fluxes. Nature 382, (1996). 29. Lynch-Stieglitz, J. et al. Atlantic meridional overturning circulation during the Last Glacial Maximum. Science 316, (2007). 30. Prange, M., Romanova, V. & Lohmann, G. The glacial thermohaline circulation: Stable or unstable? Geophys. Res. Lett. 29 (2002). 31. Knorr, G. & Lohmann, G. Southern Ocean origin for the resumption of Atlantic thermohaline circulation during deglaciation. Nature 424, (2003). 32. Sarnthein, M. et al. Changes in East Atlantic deep-water circulation over the Last 30,000 years: Eight time slice reconstructions. Paleoceanography 9, (1994). 33. Romanova, V., Prange, M. & Lohmann, G. Stability of the glacial thermohaline circulation and its dependence on the background hydrological cycle. Climate Dynamics 22, (2004). 34. Knorr, G. & Lohmann, G. Rapid transitions in the Atlantic thermohaline circulation triggered by global warming and meltwater during the last deglaciation. Geochem. Geophys. Geosyst. 8, Q12006 (2007). 35. Petit, J. R. et al. Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature 399, (1999). 36. Clark, P. U., McCabe, A. M., Mix, A. C. & Weaver, A. J. Rapid rise of sea level 19,000 years ago and its global implications. Science 304, (2004). 37. Bard, E., Rostek, F., Turon, J. L. & Gendreau, S. Hydrological impact of Heinrich events in the subtropical northeast Atlantic. Science 289, (2000). 38. Schlitzer, R. Ocean Data View, (2009). nature geoscience 9

10 supplementary information Supplementary figure annotations Figure S1. Meridional sections through the Atlantic basin show the interplay between well ventilated deep waters originating from the north and more poorly ventilated waters surrounding Antarctica. The location of TNO57-21 is identified by a white or yellow circle. Data are from GLODAP 1 and the graphics were constructed using the Ocean Data View software 38. Key: SALNT = salinity; PHSPHAT = phosphate; TCARBN = total carbon (dissolved); DELC13 = δ 13 C; DELC14 = Δ 14 C; Omega c = saturation state with respect to calcite; CONVRADCARBNAGE = conventional 14 C age. A schematic representation of the modern water mass distribution is shown on each panel ( is North Atlantic Deep Water; is Antarctic Bottom Water; is Antarctic Intermediate Water). Figure S2. Example of projection age determination. The projection age (P) is obtained by producing the decay curve for 14 C from a sample with known calendar age until it intersects with the record of atmospheric (or surface ocean) Δ 14 C 3. The orange shaded area represents the overall uncertainty (in Δ 14 C space) of the benthic sample. Errors in P are calculated as follows. A and B are the errors associated with calibration of the planktonic 14 C age to obtain its calendar age (Table S1) 4. C and D account the uncertainties both in the record of Δ 14 C and the benthic 14 C measurement error. The combined positive error on P (Fig. 2) is (A 2 +D 2 ) 0.5 while the negative error is (B 2 +C 2 ) 0.5. Figure S3. Proxies for carbonate dissolution and productivity in TNO Uppermost curve is the record of authigenic U concentration reported by Sachs and Anderson 8. The lower four curves are from this study. Coarse fraction is defined as >63μm. Shell and fragment counts were performed on the fraction >150μm and are quoted as per gram dry bulk sediment. Green shaded boxes are preservation spikes. Pink box is a period of generally poor preservation during which the fragmentation record does not agree with the records of coarse fraction or whole shells / g. This is also an interval with very low absolute numbers of fragments. %CaCO 3 data are from nature geoscience

11 supplementary information Table S1. 14 C results and calculated ventilation ages Depth Species sample 14C age error NOSAMS Cal. Age err B-P err Projection err (cm) wt (mg) (yr) (yr) Accession # (yr) (a) (yr) (yr) (yr) age (yr) (b) (yr) G. bulloides , OS , benthics , OS G. bulloides , OS , benthics , OS G. bulloides , OS , ,049 benthics , OS G. bulloides , OS , ,499 benthics , OS G. bulloides , OS , ,058 benthics , OS planktonic (c) 13, ,791 1, ,255 benthics , OS G. bulloides , OS ,869 1, ,871 benthics , OS G. bulloides , OS ,014 1, ,806 benthics , OS G. bulloides , OS ,343 1, ,127 benthics , OS G. bulloides , OS ,780 1, ,060 benthics , OS G. bulloides , OS ,451 1, ,439 benthics , OS G. bulloides , OS , ,157 benthics , OS G. bulloides , OS , ,357 benthics , OS planktonic (c) 18, ,641 1, ,449 benthics , OS G. bulloides , OS ,267 1, ,123 benthics , OS Notes: (a) Calendar ages (and 14 C ages for planktonic samples) were reported previously 4. For the two samples without planktonic 14 C ages a calendar age was determined by linear interpolation between adjacent samples with available 14 C ages. (b) The projection age was determined by combining the measured 14 C age of the benthic sample with its calendar age and back-projecting along the 14 C decay curve until intersection with the record of tropical surface ocean 14 C 3 was achieved. The surface ocean record was used instead of the atmospheric record (which is anyway derived from the ocean record over this interval) to allow better comparison with the B-P offset. The modern value of B-P is ~200yr less than the equivalent projection age due to the +200yr reservoir correction used for core location of TNO57-21 (the projection age is relative to the tropical surface ocean which has a reservoir correction of zero). (c) No planktonic age available due to sample loss during graphitisation. In order to estimate a B-P age for these samples, the calendar age was used to determine an expected 14 C age using the Marine04 calibration curve 3. nature geoscience 11

12 supplementary information Figure S1. Meridional sections through the Atlantic basin show the interplay between well ventilated deep waters originating from the north and more poorly ventilated waters surrounding Antarctica. The location of TNO57-21 is identified by a white or yellow circle. Data are from GLODAP 1 and the graphics were constructed using the Ocean Data View software 37. Key: SALNT = salinity; PHSPHAT = phosphate; TCARBN = total carbon (dissolved); DELC13 = δ 13 C; DELC14 = Δ 14 C; Omegac = saturation state with respect to calcite; CONVRADCARBNAGE = conventional 14 C age. A schematic representation of the modern water mass distribution is shown on each panel ( is North Atlantic Deep Water; is Antarctic Bottom Water; is Antarctic Intermediate Water). 12 nature geoscience

13 supplementary information Marinecal04 +/- 1σ Δ 14 C (permil) A B 100 C D 80 P Age (ka) Figure S2. Example of projection age determination. The projection age (P) is obtained by producing the decay curve for 14 C from a sample with known calendar age until it intersects with the record of atmospheric (or surface ocean) Δ 14 C 3. The orange shaded area represents the overall uncertainty (in Δ 14 C space) of the benthic sample. Errors in P are calculated as follows. A and B are the errors associated with calibration of the planktonic 14 C age to obtain its calendar age (Table S1) 4. C and D account the uncertainties both in the record of Δ 14 C and the benthic 14 C measurement error. The combined positive error on P (Fig. 2) is (A 2 +D 2 ) 0.5 while the negative error is (B 2 +C 2 ) 0.5. nature geoscience 13

14 supplementary information Age (ka) [U auth ] (ppm) whole shells / g % fragmentation Age (ka) % coarse fraction %CaCO 3 fragments / g Figure S3. Proxies for carbonate dissolution and productivity in TNO Uppermost curve is the record of authigenic U concentration reported by Sachs and Anderson 8. The lower four curves are from this study. Coarse fraction is defined as >63μm. Shell and fragment counts were performed on the fraction >150μm and are quoted as per gram dry bulk sediment. Green shaded boxes are preservation spikes. Pink box is a period of generally poor preservation during which the fragmentation record does not agree with the records of coarse fraction or whole shells / g. This is also an interval with very low absolute numbers of fragments. %CaCO3 data are from nature geoscience