Sudden spreading of corrosive bottom water during the Palaeocene Eocene Thermal Maximum

Transcription

1 SUPPLEMENTARY INFORMATION DOI: /NGEO2430 Sudden spreading of corrosive bottom water during the Palaeocene Eocene Thermal Maximum 4 Kaitlin Alexander, Katrin J. Meissner, Timothy J. Bralower 5 1 High-Alkalinity Scenarios 6 7 In order to realistically simulate the effects of Atlantic circulation on sediment dissolution, it is necessary to consider a wide range of background 8 seawater chemistries. The UVic ESCM does not simulate individual ion concentrations, but it does prognostically calculate alkalinity which depends directly on the carbonate and bicarbonate ions. Global mean alkalinity (over all depths) is specified during model spinup, to which the sediments and the carbon cycle will adjust. Global alkalinity is free to evolve during transient simulations following the spinup Figure 1 shows maximum %CaCO 3 in the pore layer (the upper layer of the sediment column, which interacts directly with the ocean) during the first 5000 years of four different simulations. Each simulation was initialised with 1680 ppm atmospheric CO 2 and forced with a 7000 GtC carbon release NATURE GEOSCIENCE 1

2 as in Meissner et al. [1], but the initial global mean alkalinity varies from the present-day value (ALK 1) to twice the present day value (ALK 2) as summarized in Table 1. Each panel is overlain with coloured points showing estimated maximum %CaCO 3 during the PETM measured at various ODP and DSDP drilling sites [2, 3, 4, 5]. ALK 1 and ALK 1.2 (Figure 1a, 1b) have unrealistically low %CaCO 3 in nearly all locations. ALK 1.5 (Figure 1c) generally agrees with CaCO 3 data except at the Southern Ocean Sites 690 and 738 as well as the Shatsky Rise Sites , where simulated %CaCO 3 is too low. ALK 2 (Figure 1d) agrees better with data in these regions, but simulated %CaCO 3 is too high throughout the North Atlantic, Caribbean, and Clipperton-Clarion Fracture Zones Figure 2 shows percent dissolution of sedimentary CaCO 3 during the same four simulations, overlain with percent dissolution estimates from CaCO 3 data at each drilling site. The ALK 1, ALK 1.2, and ALK 1.5 simulations (Figures 2a to 2c) all agree reasonably well with estimates from the sediment record. However, the ALK 2 simulation (Figure 2d) has unrealistically low dissolution in nearly all locations. Taking the data from Figures 1 and 2 into account, the ALK 1.5 simulation exhibits the best agreement with data. For this reason, it was selected for analysis in the main text A simulation with initial global mean alkalinity between ALK 1.5 and ALK 2 would likely exhibit slightly better agreement with sedimentary esti- mates of maximum %CaCO 3. It would also simulate somewhat lower percent 2

3 dissolution worldwide. In particular, such a simulation would agree with ob- servations of less severe chemical erosion at Site 690 in the Southern Ocean, which is relatively well-preserved [3] Initial Circulation Reconstructions of late Paleocene topography, bathymetry, and wind fields are identical to the configuration of Meissner et al. [1]; continental arrangement can be seen in Figure 3 of the main text. Several differences from modern geography have a significant impact on simulated ocean circulation. First, the Bering Strait is closed, preventing the flow of relatively fresh Arctic water to the North Pacific. Consequently, sea surface salinity in the North Pacific is sufficient to allow deep water formation north of 60 N. This overturning is the main source of deep water in the Northern Hemisphere. The absence of the Bering Strait also means that the Arctic Ocean is nearly closed: the Kilda Basin east of Greenland is its only connection to the global ocean. Excess freshwater from high precipitation and river runoff is essentially trapped in the Arctic Ocean, making the basin stratified and therefore oxygen-depleted at depth [1]. The freshwater also dilutes alkalinity, making the Arctic Ocean highly corrosive. The North Sea, which is fed by the Arctic Ocean through the Kilda Basin, is also low in both alkalinity and oxygen. These conditions agree with late Paleocene benthic foraminiferal assemblages 3

4 60 from the North Sea [6, 7] The Tethys Ocean connects the North Atlantic to the Indian Ocean. Tethys surface waters are warm and saline, forming intermediate water which sinks to m depth. The Panama Strait is open, which allows surface and intermediate water to flow from the North Atlantic into the tropical eastern Pacific A land bridge exists between Australia and Antarctica, which cuts off the Antarctic Circumpolar Current and instead creates a subpolar gyre southwest of Australia, in agreement with Huber et al. [8]. The Southern Ocean is the dominant source of deep water formation in all simulations. Overturning is concentrated in the Ross and Weddell Seas, forming bottom water (> 3000 m depth) which flows northward through the Atlantic and Pacific basins. A significant mass of deep water also forms around the southern tip of South America, as far north as 50 S A small amount of overturning occurs in the North Atlantic, where cold, relatively fresh, low-alkalinity water from the Arctic Ocean mixes with warm 76 and saline water from the Tethys Ocean. The mixing of these two water masses can be seen in Figure 3a, which plots the sum of Arctic and Tethys dye tracers in the pre-petm equilibrium simulation, vertically averaged above 300 m depth. The concentration of these chemically inert dye tracers is held fixed at 1 in the surface layer of the given basin, and at 0 in the surface layer elsewhere; below the surface, the dye tracers travel freely through advection 4

5 82 and diffusion. Their resulting concentrations reveal major patterns of ocean 83 circulation. Water from the Arctic Ocean flows through the Kilda Basin to the North Atlantic, where it travels southwest along the coast of North America. At the same time, water from the Tethys Ocean enters the North Atlantic and follows the subtropical gyre westward across the basin, then northeast through the Gulf of Mexico. The two water masses meet near the coast of North America, where a local maximum in the dye tracer sum can be seen. In this location, stratification is weak enough to allow deep water formation The salinity contribution from the Tethys water mass is necessary for convection; however, mixing with the corrosive Arctic water mass reduces the alkalinity of the deep water. This source of corrosive water to the deep North Atlantic can be seen in Figure 3b, which plots the vertical minimum 95 of alkalinity during the equilibrium simulation. Along the coast of North America, at precisely the location where the Arctic and Tethys water masses meet, several horizontal grid cells have a particularly low alkalinity minimum. These are the convective cells, which feed the deep North Atlantic with corrosive water Figure 4a shows the global meridional overturning streamfunction (zonally integrated meridional transport) prior to the carbon release; overturning cells from the Southern Ocean and the North Pacific are apparent. Overturning in the North Atlantic is not visible in this figure, as it is dominated 5

6 104 by North Pacific overturning at similar latitudes Changes in Circulation during the PETM Following the carbon release, deep water formation in the North Atlantic weakens during Phases 2 and 3 ( years), then strengthens during Phase 4 ( years). This process is evident in timeseries of density anomalies averaged over the deep basin (Figure 5a). Following the carbon release, density decreases gradually until Phase 4 begins. A critical value is reached around 0.8 kg m 3 below the initial conditions: at this point, the vertical 112 density gradient destabilizes the water column. For the remainder of the simulation, denser surface and intermediate waters are transported to the deep basin. By linearizing the seawater equation of state [9] with respect to temperature and salinity, we can conclude that North Atlantic overturning in Phase 4 is driven by deep ocean warming in our simulations Figure 5b shows the Atlantic Meridional Overturning Circulation (MOC) index, i.e. the vertical maximum of the zonally integrated meridional transport, evaluated at 30 N. This timeseries agrees with density anomalies, showing 4000 years of decreased activity ( 4 Sv during Phases 2-3) following the carbon release while the deep ocean warms. This is followed by a strengthening of overturning beyond initial conditions ( 7 Sv), reaching 9 Sv at the end of the simulation. 6

7 Figure 4b shows the global meridional overturning streamfunction 400 years after the carbon release (Phase 2). Since Northern Hemisphere deep water formation is dominated by the Pacific, the relatively small change in North Atlantic overturning is not visible in this plot. However, the Southern Ocean clearly shows a decrease in overturning due to surface warming. This convection cell gradually recovers until it reaches its initial state around 8000 years after the carbon release (Phase 4; Figure 4c) Atlantic Equatorial Sill The existence of corrosive bottom water in the North Atlantic depends on the equatorial sill situated on the Atlantic seafloor. This sill blocks horizontal flow and prevents the corrosive bottom water from being diluted by other water masses. There are at least four independent reconstructions of Paleogene bathymetry used in published climate model simulations, and an equatorial sill in the Atlantic is a feature of all four Huber et al. [10] used global isochron maps [11] to estimate the location of midocean ridges during the early Paleogene. These midocean ridges formed the basis for a tectonic reconstruction of early Paleogene bathymetry. The resulting equatorial sill in the Atlantic has a maximum depth between 2500 and 3500 m. Bice et al. [12] reconstructed early Eocene bathymetry by applying the age-depth relationship of Parsons & Sclater [13] to digitized 7

8 144 seafloor magnetic lineations. The equatorial sill in this reconstruction is somewhat shallower, with a maximum depth between 2000 and 3000 m [14]. Next, Tindall et al. [15] used the methods of Markwick & Valdes [16] to reconstruct early Eocene bathymetry. These GIS-based methods incorporate several sources of information, including rotation of present-day bathymetry, the age-depth relationships of Parsons & Sclater [13] andparsons & McKenzie [17], and an isochron model. The Atlantic equatorial sill as reconstructed by Tindall et al. has a maximum depth between 3600 and 4300 m [18]. Finally, Herold et al. [19] revised the early Eocene bathymetry of Müller et al. [20], which was reconstructed by applying the age-depth relationship of Stein & Stein [21] to global isochron maps. The higher resolution of this reconstruction yields a more complex shape of the Atlantic equatorial sill; nonetheless, there is a complete blockage of flow below m depth, with much of the sill around 3000 m depth In summary, an equatorial sill between 2000 and 4500 m depth in the Atlantic Ocean is consistently present in reconstructions of Paleogene bathymetry, developed using several different methods. We are therefore confident that this bathymetric feature is not an artifact of any particular reconstruction. 8

9 Model-Data Comparison of Dissolution at ODP/DSDP Sites In this section we expand upon the comparison of model output and sediment data in the main text. Figure 6 compares model estimates of CaCO 3 dissolution (red points) with those derived directly from the sedimentary 167 record (blue points) at the indicated ODP and DSDP drilling sites. As discussed by Meissner et al. [1], the reconstructed bathymetry used by the model differs from many of the paleodepths inferred by benthic foraminiferal assemblages. Consequently, we have also presented model output from points nearby to the drilling sites (within 5 longitude and 5 latitude) that lie at a more comparable depth (magenta points). All values for longitude, latitude, depth, and percent dissolution are shown in Table The purpose of Figure 6 is to illustrate the impact of differences between model bathymetry and benthic foraminiferal site depth estimates. Sediment dissolution is highly depth-dependent, since calcite saturation state gener- 177 ally decreases with depth. Additionally, many ODP and DSDP sites are situated on continental slopes. This means that small differences in a site s paleolatitude and paleolongitude (reconstructed using plate rotation models) lead to large differences in seafloor depth in reconstructions of Paleogene bathymetry. However, site paleodepth is also constrained by the species of benthic foraminifera found in the given site s sediment record. 9

10 One of the main discrepancies exists between estimates of percent dissolution based on data from Sites 401 and 549, which lie near one another in the North Atlantic. As discussed in the main text, the percent dissolution estimate at Site 401 is likely low, due to the low resolution ( 15 cm sample spacing) of the CaCO 3 data. Two measurements at and mbsf miss the minimum value as suggested by the maximum in continuous Fe-XRF values at mbsf [22]. However, Ca-XRF data do not approach zero and thus suggest lower percent dissolution than at Site 549. One explanation for this discrepancy is that the peak of Site 401, including the interval of maximum dissolution, has been obscured by a combination of bioturbation and chemical erosion, in a similar fashion to Site 1209 at Shatsky Rise [23]. A second explanation is that the peak of the carbon isotope excursion (CIE) and the interval of maximum dissolution were not recovered at the rotary-cored Site 401. Core 14, which contains the onset of the PETM, has only 74% recovery. There are several minor breaks in lithology between 95 cm and 102 cm in Section 3 of the core, the interval which includes the base of the CIE and the maximum in Fe-XRF values. This raises the possibility that a centimetre or two of sediment, including the peak in dissolution, was unrecovered in the key interval. These explanations are supported by the relatively low magnitude of the CIE at Site 401 compared to other PETM sections. For example, the benthic foraminifer N. truempyi suggests a CIE magnitude of 1.7 per mil at Site 401 [22] compared to 2.5 per mil at Site 690 [24, 25]. Similarly, the surface-dwelling foraminifer Acarinina suggests 10

11 a CIE magnitude of 3 per mil at Site 401 [22] and 4 per mil at Site 690 [24, 25] At Site 1051 in the western North Atlantic, modelled dissolution is significantly higher than the 15% dissolution calculated from the sedimentary record. However, Site 1051 likely experienced sediment failure (downslope transport) during the PETM [26], erasing the peak of the event and resulting in artificially low dissolution estimates With the exception of the incomplete section at Site 1051, model simula- tions underestimate the amount of dissolution observed at sites in the North 215 Atlantic. We note that these drilling sites are shallower than the highly corrosive water in our simulations. Consequently, this discrepancy could be due to uncertainties in paleobathymetry. For example, the equatorial sill, which is at 3000 m depth in the model bathymetry, splits the Atlantic into two regions. Below the sill, the highly corrosive bottom water leads to near-total dissolution; above the sill, horizontal mixing with less corrosive water masses means that dissolution is significantly lower. If the equatorial sill was instead at 2000 m depth, which is not unreasonable (see section 4), simulated dissolution at Sites 401, 549, and 1001 would likely be more in line with estimates from the sedimentary record. Furthermore, if the sill underlying the Panama Strait was deeper than in the reconstruction used here, the flow of intermediate water from the North Atlantic through the Caribbean and to the tropical eastern Pacific would be stronger. A larger volume of 11

12 corrosive North Atlantic bottom water would spill over the Panama sill and cause more intense dissolution at Caribbean Site 999 and Clipperton-Clarion Fracture Zones Sites 1220 and 1221, bringing simulated values more in line with the sedimentary record Model estimates of dissolution at Walvis Ridge are similar to those estimated from the sedimentary record. However, the model overestimates dissolution at Site 690 in the Southern Ocean, which is the first site the corrosive bottom water reaches after Walvis Ridge. This discrepancy is potentially due to an underestimation of the initial global mean alkalinity (see Section 1). There were also likely small-scale features on the Paleogene seafloor which channelled bottom water in the Southern Ocean. These features, which might have sheltered Site 690 from the corrosive bottom water, are not resolved in our low-resolution reconstruction of bathymetry At Shatsky Rise and possibly also Site 738 in the Southern Ocean, calculated dissolution is likely artificially low due to bioturbation, which mixed the CaCO 3 -depleted sediment of the PETM with underlying sediments higher in CaCO 3 [3]. We cannot make a detailed model-data comparison at Shatsky Rise, since there are no nearby points in the model with a comparable depth to the Shatsky Rise sites. 12

13 247 6 Alternate Scenarios The main text analysed a simulation with 1680 ppm background CO 2, an instantaneous carbon release of 7000 GtC, and an initial global mean alkalinity of mol/m 3 (1.5 times the present-day value). However, similar effects on Atlantic circulation can be seen when varying any of these three parameters. Alternate scenarios are shown in Figure 7, where timeseries of the Atlantic MOC index at 30 N (a) and sedimentary CaCO 3 dissolution averaged over the Walvis Ridge region (b) are plotted The left column of Figure 7 shows simulations initialised with 1680 ppm CO 2 followed by a 7000 GtC carbon release, but with initial global mean alkalinity ranging from the present-day value (ALK 1) to twice the present-day value (ALK 2). The two lower-alkalinity scenarios exhibit temporary jumps in MOC to a new and stronger state ( 11 Sv), which correspond to sharp spikes in dissolution at Walvis Ridge. Conversely, the two higher-alkalinity scenarios show a gradual, sustained rise in MOC once it surpasses the initial conditions of 7 Sv around 4000 years after the carbon release. As a result, the increases in dissolution seen at Walvis Ridge are more gradual than in the lower-alkalinity scenarios. These timeseries suggest that higher values for initial global mean alkalinity lead to less dramatic changes in Atlantic circulation. This relationship can be explained by the fact that higher alkalinity enhances the ability of the ocean to absorb atmospheric CO 2. As a result, a smaller proportion of the carbon release remains in the atmosphere, leading 13

14 to less warming compared to lower-alkalinity scenarios. Since the simulated changes in Atlantic circulation appear to be temperature-driven (Figure 5), less dramatic changes in MOC will occur The centre column of Figure 7 shows simulations initialised with 1680 ppm CO 2 and global mean alkalinity 1.5 times the present-day value, but with the magnitude of the carbon release ranging from 3000 GtC to GtC. It is clear that larger carbon releases lead to more dramatic changes in MOC (in fact, a carbon release of GtC triggers a temporary jump to the elevated MOC state discussed previously) and more intense spikes in dissolution at Walvis Ridge. Again, this is because larger carbon releases lead to more intense warming The right column of Figure 7 shows simulations initialised with presentday alkalinity and forced with a 7000 GtC carbon release, but with initial CO 2 levels varying from 840 ppm to 2520 ppm. Higher initial CO 2 levels lead to more dramatic changes in MOC: both the 1680 ppm and the 2520 ppm scenarios jump to the elevated MOC state, but it is sustained for much longer in the 2520 ppm scenario. This behaviour indicates that simulated changes in Atlantic circulation are unrelated to the magnitude of temperature change; otherwise we would expect the opposite trend, since a given release of carbon will cause more warming in a low-co 2 atmosphere than in a high-co 2 atmosphere. Rather, the circulation changes seem to be related to absolute temperature. Spikes in dissolution at Walvis Ridge following enhanced proto- 14

15 NADW formation are short-lived for all three scenarios, since the relatively low alkalinity leads to very little initial sediment available to dissolve Simulations initialised with present-day alkalinity and forced with a gradual carbon release of 1 GtC/year for 4500 years, as described by Meissner et al. [1], were also performed. Each simulation exhibited the same changes in Atlantic circulation and sediment dissolution that occurred when 4500 GtC was instead released instantaneously, but these changes occurred approximately 2000 years later in the gradual simulations (not shown). 15

16 299 References [1] Meissner, K. J., et al. The Paleocene-Eocene Thermal Maximum: How much carbon is enough? Paleoceanography 29, (2014) [2] Pälike, C., Delaney, M. L., & Zachos, J. C. Deep-sea redox across the Paleocene-Eocene thermal maximum. Geochemistry, Geophysics, Geosystems 15, (2014) [3] Bralower, T. J., Meissner, K. J., Alexander, K., & Thomas, D. J. The dynamics of global change at the Paleocene-Eocene Thermal Maximum: A data-model comparison. Geochemistry, Geophysics, Geosystems 15, (2014) [4] Panchuk, K. M. Investigating the Paleocene-Eocene Carbon Cycle Per- turbation: An Earth System Model Approach. Pennsylvania State Uni- versity, Graduate School, Department of Geosciences (2007) [5] Dunkley Jones, T., et al. Climate model and proxy data constraints on ocean warming across the Paleocene-Eocene Thermal Maximum. Earth- Science Reviews 125, (2013) [6] Thomas, E. Biogeography of the Late Paleocene Benthic Foraminiferal Extinction. Division III Faculty Publications 300, (1998). 16

17 [7] Kender, S., et al. Marine and terrestrial environmental changes in NW Europe preceding carbon release at the Paleocene-Eocene transition. Earth and Planetary Science Letters , (2012) [8] Huber, M., et al. Eocene circulation of the Southern Ocean: Was Antarc- tica kept warm by subtropical waters? Paleoceanography 19, PA4026 (2004) [9] UNESCO. Tenth report of the joint panel on oceanographic tables and standards. Unesco Technical Papers in Marine Science 36, 1-25 (1981) [10] Huber, M., Sloan, L. C., & Shellito, C. Early Paleogene oceans and climate: A fully coupled modeling approach using the NCAR CCSM. Geological Society of America Special Papers 369, (2003) [11] Royer, J. Y., et al. A global isochron chart. University of Texas Institute for Geophysics Technical Report 117, 1-38 (1992) [12] Bice, K. L., Barron, E. J., & Peterson, W. H. Reconstruction of realistic Early Eocene paleobathymetry and ocean GCM sensitivity to specified basin configuration, in Tectonic Boundary Conditions for Climate Reconstructions (Oxford University Press, Oxford, 1998) [13] Parsons, B. & Sclater, J. G. An analysis of the variation of ocean floor bathymetry and heat flow with age. Journal of Geophysical Research 82, (1977). 17

18 bilized methane hydrate at the Paleocene/Eocene boundary? ceanography 17, 1018 (2002). [14] Bice, K. L. & Marotzke, J. Could changing ocean circulation have desta- Paleo [15] Tindall, J., et al. Modelling the oxygen isotope distribution of ancient seawater using a coupled ocean-atmosphere GCM: Implications for reconstructing early Eocene climate. Earth and Planetary Science Letters 292, (2010) [16] Markwick, P. J., & Valdes, P. J. Palaeo-digital elevation models for use as boundary conditions in coupled ocean-atmosphere GCM experiments: a Maastrichtian (late Cretaceous) example. Palaeogeography, Palaeoclimatology, Palaeoecology 213, (2004) [17] Parsons, B., & McKenzie, D. Mantle convection and the thermal struc- ture of the plates. Journal of Geophysical Research: Solid Earth 83, (1978) [18] Lunt, D. J., et al. A model for orbital pacing of methane hydrate desta- bilization during the Palaeogene. Nature Geoscience 4, (2011) [19] Herold, N., et al. A suite of early Eocene ( 55 Ma) climate model bound- ary conditions. Geoscientific Model Development 7, (2014) [20] Müller, R. D., Sdrolias, M., Gaina, C., Steinberger, B., & Heine, C. Long-Term Sea-Level Fluctuations Driven by Ocean Basin Dynamics. Science 319, (2008). 18

19 [21] Stein, C. A., & Stein, S. A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359, (1992) [22] Bornemann, A., et al. Persistent environmental change after the Paleocene-Eocene Thermal Maximum in the eastern North Atlantic. Earth and Planetary Science Letters 394, (2014) [23] Bralower, T. J., et al. Impact of dissolution on the sedimentary record of the Paleocene-Eocene thermal maximum. Earth and Planetary Science Letters 401, (2014) [24] Thomas, D. J., Zachos, J. C., Bralower, T. J., Thomas, E., and Bohaty, S. Warming the fuel for the fire: Evidence for the thermal dissociation of methane hydrate during the Paleocene-Eocene thermal maximum. Geology 30, (2002) [25] Kennett, J. P. & Stott, L. D. Abrupt deep-sea warming, palaeoceano- graphic changes and benthic extinctions at the end of the Palaeocene. Nature 353, (1991) [26] Katz, M. E., Pak, D. K., Dickens, G. R., & Miller, K. G. The Source and Fate of Massive Carbon Input During the Latest Paleocene Thermal Maximum. Science 286, (1999) [27] Winguth, A. M. E., Thomas, E., and Winguth, C. Global decline in ocean ventilation, oxygenation, and productivity during the Paleocene- 19

20 Eocene Thermal Maximum: Implications for the benthic extinction. Ge- ology 40, (2012). 20

21 Simulation Global mean alkalinity Value ALK 1 present-day mol/m 3 ALK times present-day mol/m 3 ALK times present-day mol/m 3 ALK 2 2 times present-day mol/m 3 Table 1: Initial global mean alkalinity for simulations shown in Figures 1, 2, and 7. 21

22 22 Site Paleo- Paleo- Paleo- % dissl Depth % dissl Nearby Nearby Depth % dissl longitude latitude depth (data) (model) (model) longitude latitude (model) (model) W N 1900 m m W 42 N 1683 m W N 1000 m m W 44 N 1683 m W 28.9 N 1500 m m W 33 N 1378 m W 10.8 N 2000 m m W 13 N 2018 m W 13.6 N 2300 m m W 3.08 S 2900 m m W 1.23 S 3200 m m W S 3600 m m W S 1500 m m W 37 S 1378 m W S 2600 m m W 35 S 2383 m W S 3200 m m W 39 S 3203 m E 65.4 S 2000 m m S 1683 m E S 1800 m m E 66 S 2018 m W 22 N 2500 m m W N 2100 m m W N 2400 m m W 21.8 N 2200 m m Table 2: Location and percent dissolution data used in Figure 6. Paleolongitude and paleolatitude are given by Dunkley Jones et al. [5] combined with modern latitude-longitude offsets between nearby sites. Paleodepths are given by Dunkley Jones et al. [5], Panchuk [4], and Winguth et al. [27]. CaCO 3 measurements used to calculate estimates of percent dissolution from data are compiled by Panchuk [4] as well as Pälike et al. [2] (Site 401) and Bralower et al. [3] (Site 690). In many locations, particularly on continental slopes, the seafloor depth in the model was significantly deeper than paleodepth estimates for the given site. If a nearby location (within 5 latitude and longitude) with a more comparable depth was available, we recalculated the percent dissolution as simulated by the model at that location.

23 Figure 1: Maximum %CaCO 3 in the pore layer during the first 5000 years of each 1680 ppm GtC simulation, with global mean alkalinity varying from the present-day value (2.429 mol/m 3 ) to twice the present-day value (4.858 mol/m 3 ). Coloured points overlaying the model output show calculated values for maximum %CaCO 3 during the PETM at the indicated ODP and DSDP drilling sites [2, 3, 4, 5]. 23

24 401 Figure 2: Percent dissolution of sedimentary CaCO 3 during each 1680 ppm GtC simulation, with global mean alkalinity varying from the presentday value (2.429 mol/m 3 ) to twice the present-day value (4.858 mol/m 3 ). Coloured points overlaying the model output show percent dissolution calculated from sedimentary data at the indicated ODP and DSDP drilling sites [2, 3, 4, 5]. 24

25 Kilda Basin North Sea Figure 3: a) The sum of Arctic and Tethys Ocean dye tracers during the pre- PETM equilibrium simulation, averaged above 300 m. b) Vertical minimum of alkalinity during the equilibrium simulation. The particularly corrosive cells off the coast of North America are the convective cells. Figure 4: Simulated global meridional overturning streamfunction in Sv (10 6 m 3 /s) at three timesteps: a) Phase 1, prior to the carbon release; b) Phase 2, 400 years after the carbon release; c) Phase 4, 8000 years after the carbon release. 25

26 a) Deep North Atlantic Density kg m ρ ρ temp ρ sal b) Atlantic MOC at 30 N 8 Sv years since carbon release Figure 5: a) Timeseries of density anomalies ( ρ) averaged over the North Atlantic below 3000 m depth. Also shown are contributions of temperature ( ρ temp ) and salinity ( ρ sal ) to ρ, which were computed by linearizing the seawater equation of state [9] with respect to temperature and salinity. Vertical lines at years 500 and 4000 show the boundaries of Phases 2-4. b) Timeseries of the Atlantic MOC index (vertical maximum of the zonally integrated meridional transport) at 30 N during the same simulation. 26

27 North Atlantic Caribbean Clipperton Clarion Fracture Zones depth (m) Walvis Ridge Southern Ocean Shatsky Rise depth (m) Figure 6: Percent dissolution of sedimentary CaCO 3 at the indicated ODP and DSDP drilling sites. Blue values show sedimentary data as in Figure 2, with depth estimates from Dunkley Jones et al. [5], Panchuk [4], and Winguth et al. [27]. Red values show model results interpolated to the site s paleolocation as shown in Figure 3c of the main text, with depth estimates taken as the depth of the model seafloor at that location. Magenta values show model results at nearby locations (within 5 longitude and 5 latitude of the drilling site) where the depth of the model seafloor is more comparable to the sedimentary depth estimates. All values are also listed in Table 2. 27

28 a) Atlantic MOC at 30 N 12 Varying Alkalinity 12 Varying Carbon Release 12 Varying Initial CO Sv years since carbon release b) Sediment CaCO3 Dissolution (Walvis Ridge) 10 9 mol m 2 s Varying Alkalinity ALK 1 ALK 1.2 ALK 1.5 ALK Varying Carbon Release GtC 4500 GtC GtC GtC years since carbon release Varying Initial CO2 840 ppm ppm 2520 ppm Figure 7: Timeseries of a) the Atlantic MOC index at 30 N and b) sediment dissolution averaged horizontally over the Walvis Ridge region shown in Figure 3b of the main text. The left column shows simulations initialised with 1680 ppm CO 2 followed by a 7000 GtC carbon release, with initial global mean alkalinity varying from the present-day value (2.429 mol/m 3 ; denoted ALK 1) to twice the present-day value (4.858 mol/m 3 ; denoted ALK 2). The centre column shows simulations initialised with 1680 ppm CO 2 and global mean alkalinity 1.5 times the present-day value, but with the carbon release varying from 3000 GtC to GtC. The right column shows simulations initialised with present-day alkalinity and forced with a 7000 GtC carbon release, but with initial CO 2 levels varying from 840 ppm to 2520 ppm. 28