Supplementary Online Material
Age control
Radiocarbon ages for the Dry Tortugas and Great Bahama Bank cores are based
on > 5 mg samples of Globigerinoides ruber from the > 250 m size fraction. Several
core tops have >1 fraction modern radiocarbon, indicating that deposition occurred
primarily after the beginning of nuclear weapons testing in the mid-20th century (62MC,
118MC, 125MC, 133GGC, 134MC) (Tables 1s & 2s). Others have 14C core top ages less
than the 400-year surface water reservoir age for this area (3MC, 11MC, W167-79GGC,
94MC), indicating the presence of bomb radiocarbon, but to a lesser extent than those
with >1 fraction modern 14C. The only KNR166-2 sites without modern or near-modern
core top material are those shallower than 400 m water depth on the Dry Tortugas side of
the Florida Current. We dated a total of 10 separate cores from 200 to 400 m in this
region and none had core top ages < 900 14C years. Of these cores, three display a
monotonic increase in age with depth (16MC, 49GGC, 50MC). Core 16MC (250 m
water depth) has the most recent material, with a core top age of 960 14C years. All raw
radiocarbon ages were converted to calendar ages using CALIB 5.0 1 assuming a surface
ocean reservoir age of 400 years. Age models for each core were created by linear
interpolation between 14C control points.
The sedimentation rates for the Dry Tortugas and Great Bahama Bank sites are
uniformly high – all but one of the cores has a sedimentation rate in excess of 20 cm/kyr
(The exception is 11MC, which has an accumulation rate of 11 cm/kyr.) On either side
of the Florida Current, mean sedimentation rates increase with decreasing water depth.
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The highest accumulation rates are on the Great Bahama Bank, with two cores (93GGC,
and 117GGC) in excess of 100 cm/kyr.
Sampling strategy
Stable isotope analyses for the Florida Straits cores were based primarily on
Cibicidoides species from the >250 m size fraction (Figures 1s and 2s). At the deeper
Dry Tortugas and Great Bahama Bank sites, we used Planulina ariminensis when
necessary, also from the >250 m size fraction. These species follow the fractionation
for inorganic calcite precipitation determined by Kim and O’Neil 2, where the 18Ocalcite18Oseawater difference follows a thermodynamic slope of ~0.2 per mil/C 3. We ran an
average of four 18O analyses at each stratigraphic level, each based on an individual
benthic foraminifer test. Sonication of foraminifera to remove attached detrital material
was performed as necessary and each foraminiferal test was crushed prior to analysis to
ensure complete acidification. Stable isotope analyses were run on a Finnigan MAT 253
coupled to a Kiel III carbonate device. Calibration to VPDB scale was made using NBS19 (18O = -2.20‰ and 13C = 1.95‰). Long-term reproducibility (1) of NBS-19
(n=461) for this mass spectrometer system is 0.08‰ for 18O and 0.04‰ for 13C.
Density calibration
To quantify the relationship between foraminiferal 18O (18Oc) and seawater t,
we first estimated seawater 18O (18Ow) for CTD data taken at the core sites in January
2002 using the 18Ow-S relationship from tropical and subtropical thermocline waters in
the Atlantic (Figure 3s, top). We then calculated the predicted 18Oc of Cibicidoides and
2
Planulina precipitated in equilibrium with this water [18Oc = (18Ow – 0.27) – 0.21*T +
3.38] (3). Finally, we determined the t values for the CTD data, and created a regression
between t and 18Oc (Figure 3s, bottom). The resulting polynomial fit (t = 25.99 +
1.07*18Oc – 0.17*18Oc2) was used to estimate past changes in seawater density.
Comparison of core top foraminiferal and CTD density values from Dry Tortugas
and Great Bahama Bank indicates the stable oxygen isotopic composition of the benthic
foraminifera used in this study reliably record the density structure of the Florida Current
(Figure 4s). The foraminifera clearly capture the increase in density with depth at each
hydrographic station, the sharp pycnocline below the surface mixed layer, and the
gradient in density between the Great Bahama Bank and Dry Tortugas. Offsets between
the CTD and foraminifera t data are most likely a function of short-term variability in
Florida Current density not captured by the foraminiferal records. Florida Current
transport varies by up to 8 Sv on weekly to monthly timescales4 and 2-4 Sv on annual
to interannual time scales5. A portion of this variability is baroclinic4 and should affect
Florida Current density structure. It is therefore not surprising that the density values
from individual CTD casts are slightly different than those inferred from core top
foraminifera, which represent average conditions for the last ~50 years.
The level of no motion at 850 m water depth (see next section) is deeper than our
core sites at 750 m near Dry Tortugas and 700 m on Great Bahama Bank. As a result, it
was necessary to estimate density between these core depths and 850 m using CTD data.
At Dry Tortugas, we estimated density at 850 m by adding 0.05 t units to the values
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from 750 m water depth. The January 2002 CTD data indicate t increases by 0.05 per
100 m in the 600-800 m depth range (Figure 4s), similar to that observed from 800-900 m
at Dry Tortugas6. For Great Bahama Bank, we added 0.15 t units to the values from 700
m water depth to estimate density at 850 m. Density increases by 0.1 t units per 100 m
in the 600-800 m depth range (Figure 4s), very similar to the trend from 800-900 m near
Havana, Cuba6.
Transport calculation
We estimated Florida Current transport using the geopotential method. This
method utilizes the specific volume anomaly (), which requires independent knowledge
of temperature and salinity. Since we lack this information, we instead use the
thermosteric anomaly (s,t), which is the primary component of  in the upper 1000 m of
the ocean and can be calculated directly from t 7. We also approximate pressure (p) at
any given depth (z) using p = z*104 Pa 7. The error compared to using the hydrostatic
equation is minimal (<1% using January 2002 data) since the approximation is similar at
each station and largely cancels when geopotential anomalies are differenced.
The level of no motion is both the most important and most poorly defined
variable used in the geostrophic technique. Typically, a level of no motion is assumed
near the bottom of flow, but the exact depth at which this occurs is difficult to constrain
even when independent velocity measurements are available. Today, current meters
show that near-zero velocities between Cay Sal Bank and Florida occur between 800 m
and 1000 m 8. Offshore of Key West, the level of no motion is deeper (~ 1000 m) but all
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velocities below 600 m are less than 5 cm/s 9. In the much shallower Santaren Channel,
near-zero flow velocities occur at 600 m water depth 8. Given the wide range of possible
choices, we used a level of no motion (850 m) that yields a core top transport of 31 Sv,
consistent with modern instrumental estimates.
If the level of no motion changed during the last millennium, it could alter the
flow time series. For example, a shift in the level of no motion from 850 to 950 m during
the Little Ice Age, but with the same shear, would produce a total volume transport of
~31 Sv, equal to the modern value (Figure 5s). Increasing the depth-averaged flow above
850 m by ~4 cm/s has a similar effect. On a seasonal basis, the mean flow of the Florida
Current at Miami changes by 10-20 cm/s 4. In the context of the modern seasonal cycle,
shifts in the mean flow of 4 cm/s seem plausible. This is a poor analog for the
foraminifera-based reconstruction, however, given the vastly different time scales
involved and the higher bottom velocities observed near Miami4 relative to Cay Sal
Bank8. It is also possible that there was a decrease in mean flow or shoaling of the level
of no motion during the LIA, which would amplify the transport signal. At this stage,
these possibilities are unconstrained. Therefore, we assume that the level of no motion
remained near 850 m during the last millennium.
18Ow-S transport error
To calculate past changes in density, we assumed the thermocline salinity and
seawater 18O (18Ow) relationship was constant through time (Figure 1s). Yet planktonic
18Oc and Mg/Ca time series from the Dry Tortugas and Great Bahama Bank imply there
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was substantial variability in mixed layer 18Ow during the last millennium10. These
changes are most reasonably explained by invoking an increase in mean thermocline
18Ow during the LIA. Although the planktonic data do not require a change in either the
slope or intercept of the thermocline 18Ow-S regression, it is possible that the 18Ow-S
relationship changed during the last millennium.
One way to alter the thermocline 18Ow-S relationship would be to increase the
meridional sea-surface temperature gradient, which would cause a greater depletion in
18
O of precipitation at high-latitudes. If temperatures in the region where subtropical
gyre isopycnals outcrop cooled by ~1ºC during the LIA11,12 , this would decrease rainfall
18O by 0.7‰ based on the relationship between 18O of precipitation and mean annual
air temperature on land13. Since continental rainfall is generally more depleted in 18O
than oceanic rainfall, 0.7‰ is probably an overestimate of the LIA cooling effect on
mean annual 18O/16O of oceanic precipitation. This effect may be offset by a greater
temperature gradient in winter, however, when the thermocline is ventilated. As an
approximation, we assume that LIA cooling lowered the 18O of the fresh-water end
member by up to 1‰.
An isotopic depletion of the fresh-water end member results in a steeper 18Ow-S
slope (Figure 6s, top) which in turn alters the polynomial fit to the 18Oc -t data in the
range occupied by the Dry Tortugas and Bahamas benthic 18Oc values (Figure 6s,
middle). The net effect is to increase density at Dry Tortugas more than the Bahamas,
producing slightly higher transport values (~0.5 Sv) (Figure 6s, bottom). Only if the
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fresh end member decreased by 4‰ would the resulting transport bias (~ 3 Sv) approach
the observed signal for the last millennium3. We therefore believe that the reconstructed
transport is robust with respect to realistic long-term variability in the 18O of
precipitation at high latitudes.
References
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synthetic carbonates. Geochim. Cosmochim. Acta 61, 3461-3475 (1997).
Lynch-Stieglitz, J., Curry, W. B., and Slowey, N. A geostrophic estimate for the
Florida Current from the oxygen isotope composition of benthic foraminifera.
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Niiler, P. P. and Richardson, W. S. Seasonal variability of the Florida Current.
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Baringer, M. O., and Larsen, J. C. Sixteen years of Florida Current transport at
27°N. Geophys. Res. Lett. 28, 3179-3182 (2001).
Montgomery, R. B. Transport of the Florida Current off Habana. Journal of
Marine Research, IV (3), 198-220 (1941).
Pond, S., and Pickard, G. L. Introductory Dynamical Oceanography, 2nd edition,
Pergammon Press, Oxford, 329 pp. (1983).
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Waddell, E., Transport, potential vorticity, and current/temperature structure
across Northwest Providence and Santaren Channels and the Florida Current off
Cay Sal Bank. J. Geophys. Res., 100 (C5), 8561-8569 (1995).
Hamilton, P., Larsen, J. C., Leaman, K. D., Lee, T. N., and Waddell, E.
Transports through the Straits of Florida, Journal of Physical Oceanography 35,
308-322 (2005).
Lund, D. C., & Curry, W. B., Florida Current surface temperature and salinity
variability during the last millennium. Paleoceanography 21, PA2009,
doi:10.1029/2005PA001218 (2006).
Keigwin, L. D. The Little Ice age and Medieval Warm Period in the Sargassso
Sea, Science 274, 1504-1508 (1996).
Marchitto, T. M., & deMenocal, P. B. Late Holocene variability of upper North
Atlantic Deep Water temperature and salinity. Geochem. Geophys. Geosyst.
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Dansgaard, W. Stable isotopes in precipitation, Tellus XVI (4), 436-468 (1964).
Schmidt, G.A., Bigg, G. R., and Rohling, E. J. Global Seawater Oxygen-18
Database. [http://www.giss.nasa.gov/data/o18data/] (1999).
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Deuser, W. G. Seasonal variations in isotopic composition and deep-water fluxes
of the tests of perennially abundant planktonic foraminifera of the Sargasso Sea:
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(2005).
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Figure 1s
Benthic 18O for the Dry Tortugas cores at 200, 250, 450, 540, and 750 m water depth.
Each symbol represents an individual 18O analysis. The black line is the average 18O
value at each stratigraphic level. The vast majority of analyses was made using
Cibicidoides species (o), though below 450 m water depth Planulina ariminensis (+) was
also used.
Figure 2s
Benthic 18O for the Great Bahama Bank cores at 260, 440, 530, and 700 m water depth.
Each symbol represents an individual 18O analysis. The black line is the average 18O
value at each stratigraphic level. Similar to the Dry Tortugas, most analyses were made
using Cibicidoides (o), though at 700 m water depth we also used Planulina ariminensis
(+).
Figure 3s
Top) 18O of sea water (18Ow) versus salinity regressions. One is based on 18Ow -S
pairs from GEOSECS stations 29 and 37 (thin line; after Lynch-Stieglitz et al.3), and the
other is based on a broader set of open-ocean thermocline data from the North Atlantic
subtropical gyre and the tropical Atlantic (circles and thick line; Schmidt et al.14).
Bottom) Estimates of t for given values of 18Oc based on the two different 18Ow –S
relationships. Since the Lynch-Stieglitz et al.3 polynomial is based on a more limited
18Ow-S data set, we used t = 25.99 + 1.07*18Oc – 0.17*18Oc2 to estimate past changes
in seawater density.
Figure 4s
Comparison of sigma-t vs. depth based on the 18Oc of core top foraminifera (solid
symbols) with that calculated using T & S data from January 2002 CTD casts.
The G. ruber data (symbols at 0 and 50 m water depth) reflect annual average mixed
layer conditions15 and are therefore offset relative to the wintertime CTD data. If we
adjust the Great Bahama Bank G. ruber 18Oc values by +0.5 per mil to account for the
difference between annual average and wintertime SSTs (2.5ºC) 16, then the resulting
mixed layer t (24.2) is similar to that observed in January of 2002. If core top Dry
Tortugas G. ruber 18Oc is adjusted by +0.6 per mil to account for the 3ºC offset between
annual average and winter SSTs 16, then t = 25.0, indistinguishable from the mixed layer
CTD data. For the transport calculations, we used an annual average mixed layer depth
of 50 m at both locations.
Figure 5s
Velocity and transport profiles for today (thin black line), 200 yr BP (thick black line),
and 1100 yr BP (grey line). Relative velocities (Vrel) were calculated assuming a scalewidth of 100 km for the Florida Current. We define velocity to be 0 cm/s at 850 m, the
level of no motion (arrow). A +0.4 cm/s increase in the mean or barotropic flow during
the LIA (dashed line) compensates for the effect of reduced baroclinic shear, yielding a
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total transport of ~31 Sv, similar to the modern value. Moving the level of no motion to
950 m during the LIA has nearly the same effect (dotted line). (Total transport is 0.2 Sv
higher for a deeper level of no motion given the additional flow below 850 m.)
Figure 6s
Transport sensitivity to changes in the 18Ow of the freshwater end member. Top) The
original (solid) and altered (dash) thermocline 18Ow-S relationships, the latter having a 1
per mil more negative fresh end member. Middle) The effect of the alternative 18Ow-S
slope on the 18Oc-t regression. In the range of the benthic 18Oc data (horizontal bars),
the altered 18Oc-t regression (dashed line) increases t more for the Dry Tortugas than
the Bahamas, and as a result amplifies the cross-current density gradient. Bottom) The
resulting effect on transport (dashed line) is small (<0.5 Sv) relative to that based on the
original 18Ow-S relationship (solid line).
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