1
2
Evidence for increased silica leakage to the tropical Atlantic during glacial
development – Supplementary Material
3
4
5
James Griffiths, Stephen Barker, Katharine Hendry, David Thornalley, Tina van de Flierdt,
Robert Anderson, Ian Hall
6
7
Supplementary Information
8
Core location hydrography
9
Caribbean Sea surface waters (0-80 m) are nutrient-depleted and are underlain by high-
10
salinity Subtropical Under Water (SUW) between ~80-100 m. Below this, a mixture of SUW
11
and AAIW forms the main component of Caribbean thermocline waters present at depths
12
~500-900 m. The composition of intermediate waters at this site (termed Atlantic
13
Intermediate Water - AIW, at depths ~900-1900 m) is a mixture of AAIW and Upper North
14
Atlantic Deep Water (UNADW). UNADW dominates below 1900 m [Wüst, 1964; Haddad
15
and Droxler, 1996]. AAIW enters the Tobago Basin via the subthermocline North Brazil
16
Current (NBC), which contains approximately 60 ± 5% southern-sourced water [Bub and
17
Brown, 1996]. North Atlantic Central Water (NACW) enters the region between the equator
18
and 9°N and west of 44°W, and mixes with the southern-sourced water. This mixed water
19
mass makes up approximately half of the region in volume and is predominantly of a
20
southern origin [Bub and Brown, 1996].
21
The ODP core site 1063 is situated on the Bermuda Rise, an area of very high sediment
22
accumulation, mainly sourced from Canadian rivers [Laine and Hollister, 1981] (Figure 2).
23
Sediments on the north side of the Bermuda Rise are resuspended by the Gulf Stream and
24
deep recirculating gyres [Laine and Hollister, 1981]. Primary productivity is low due to the
25
oligotrophic surface waters bounded by the North Atlantic Subtropical Gyre (NASG), and
1
26
modern biogenic silica production rates are amongst the lowest in the world ocean
27
[Brzezinski and Nelson, 1995; Nelson and Brzezinski, 1997].
28
The RC24-01 core site lies in the divergence created by the boundaries of the broad,
29
westward-flowing South Equatorial Current (SEC) and the weaker, more variable eastward-
30
flowing Northern Equatorial Counter Current (NECC) [Bourles et al., 1999; Stramma and
31
Schott, 1999]. Salinity profiles demonstrate that AAIW is present as a layer in the EEA at
32
depths of 800-1000m [Schlitzer, 2000]. In the modern ocean AAIW influences the EEA after
33
flowing eastward from the NBC at 3 - 4°S of the equator and also more weakly at 1 - 2°N
34
[Suga and Talley, 1995].
35
36
Methods
37
Neodymium isotope ratios were measured on the dispersed Fe-Mn oxyhydroxide phase
38
extracted from the fine (<63 μm) fraction of the de-carbonated bulk sediment from MD99-
39
2198 sediment samples. In detail, the <63μm fraction of the sample (~4cm3 bulk) was
40
separated out by wet sieving. All carbonate was removed from the sample using buffered
41
acetic acid, until no signs of a reaction were detectable. Fe-Mn oxides were subsequently
42
extracted using a 0.02 M solution of hydroxylamine hydrochloride (HH) for 2 hours, after
43
Chester and Hughes [1967]. After drying down the solution at a high temperature, the
44
samples were redissolved in 3M HNO3- for two-stage ion chromatography. Separation of rare
45
earth elements (REE) from the sample matrix was achieved using TRU-spec resin, and
46
separation of Nd from the other REE was achieved using Ln-spec resin following standard
47
procedures. Neodymium isotopes were measured in static mode on a Nu Plasma Multi-
48
Collector Inductively-Coupled Plasma Mass Spectrometer (MC-ICP-MS) in the MAGIC
49
laboratories at Imperial College London. Mass bias correction was accomplished using a
2
50
146
Nd/144Nd ratio of 0.7219. During the course of the sample analyses (3 separate days) JNdi
51
standards yielded
52
0.000012 (2σ SD, n=14), and 0.512057 ± 0.000015 (2σ SD, n=19), respectively. All sample
53
values were normalised to the recommended JNdi
54
al., 2000] (see Table 1).
143
Nd/144Nd values of 0.512105 ± 0.000017 (2σSD, n=21) , 0.512112 ±
143
Nd/144Nd ratio of 0.512115 [Tanaka et
55
Sponge spicules were picked from the 63-215µm fraction of the previously-separated
56
coarse (>63µm) fraction of MD99-2198. The species from which the spicules came were not
57
identified as previous work had suggested that the species of sponge does not affect the
58
relationship between [Si(OH)4] and δ30Si [Hendry et al., 2010]. The sponge spicules were
59
cleaned in H2O2, and dissolved in 0.4 M NaOH at 100°C for three days. The solutions were
60
then diluted and acidified to pH~2-3. A cation exchange resin (BioRad AG50W-X12) was
61
used to quantitatively separate Si from other major ions [Georg et al., 2006]. Si isotopes were
62
measured on a Thermo Neptune Multi-Collector Inductively Coupled Plasma Mass
63
Spectrometer at Woods Hole Oceanographic Institution. Full operating conditions are
64
described in Hendry et al. [2010]. Solutions were run at least in duplicate by both standard-
65
sample bracketing and bracketing with Mg-doping, whereby samples and bracketing
66
standards were spiked with Mg standard (Inorganic Ventures), and intensity-matched for 28Si
67
and
68
corrected using a fractionation factor calculated using the measured
69
Cardinal et al. [2003] for details. Repeat analyses of an opal standard LMG08 [Hendry et al.,
70
2011] measured by standard-sample bracketing of an opal standard over a period of several
71
months yields values of -1.75 ± 0.1‰ for δ29Si and -3.43 ± 0.23‰ (2SD) for δ30Si, providing
72
the most conservative estimate of uncertainty, see Table 2.
Mg signals within 10% (typically within 5%). The
29
Si/28Si isotope ratios were
25
Mg/24Mg ratios, see
Concentrations of protactinium (231Pa), thorium (230Th, 232Th), and uranium (234U and
73
74
24
238
U) in RC24-01 were determined at Lamont-Doherty Earth Observatory (L-DEO),
3
75
according to the method of Anderson and Fleer [1982] and Fleisher and Anderson [2003].
76
Approximately 0.1g of bulk sediment was digested using HNO3-, HClO4 and HF. Anion resin
77
column chemistry was then used to separate out the Pa and U/Th fractions. The concentration
78
of each radionuclide species was then measured using a VG Axiom Multi-Collector
79
Inductively-Coupled Plasma Mass Spectrometer. The Pa isotopes were measured in ‘Flight
80
Tube Scan Mode’, and only monitored the masses 231 and 233. Precision on the 231/233
81
ratio was better than 2.5% at the 1 sigma level, for 5 replicates. See Fleisher and Anderson
82
[2003] for full operating conditions. Authigenic uranium (Uauth) was determined from the
83
concentrations of
84
[Bradtmiller et al., 2007]. To account for sedimentary redistribution, we normalised the
85
sedimentary opal content to the flux rate of
86
1984; Francois et al., 2004]. Opal flux was calculated as follows: 0.01 x
87
opal (see Table 3 for data). Opal was measured in all cores using ~100mg of bulk sediment.
88
Sedimentary opal content was measured using the wet alkaline extraction method of
89
Mortlock and Froelich [1989].
232
Th and
238
U, assuming a detrital
230
232
Th contribution of 10ppm
Th, to produce a record of opal flux [Bacon,
230
Th-norm F x %
90
91
Core age control
92
We used the North Greenland Ice Core Project (NGRIP) δ18O record (using the
93
GICC05 timescale (Andersen et al., 2007) for 0-60 ka, and a speleothem-tuned age model
94
prior to 60 ka) [Barker et al., 2011] as a tuning target for all age models used in this study.
95
Age control for MD99-2198 was initially based on a low resolution record of benthic 18O
96
(R. Zahn unpublished data; Supplementary Material) that allowed identification of MIS 4 and
97
fine-tuned by aligning a higher resolution planktonic 18O record (Globigerinoides ruber,
98
white, picked from the 250-315 μm fraction) with the NGRIP 18O record (Figure 3). This
4
99
approach assumes in-phase behaviour between millennial-scale oscillations in the tropics and
100
the northern hemisphere temperature. We suggest that this is reasonable, because these
101
regions are linked through meridional heat transport and the position of the Intertropical
102
Convergence Zone (ITCZ) (which is sensitive to changes in North Atlantic temperature)
103
[Hüls and Zahn, 2000; Peterson et al., 2000; Lea et al., 2003; Cruz et al., 2005]. It has also
104
been demonstrated that the primary control on the δ18O of rainfall over tropical South
105
America is the amount of precipitation, which is modulated by the position of the ITCZ
106
[Vuille et al., 2003; Cruz et al., 2005]. The uncertainty associated with individual control
107
points in our age model (with respect to the Greenland record to which it is tuned) depends
108
on our ability to identify the correct transitions for tuning and the precision with which tuning
109
can be performed (which is dependent on the duration of the transition). However, much of
110
our discussion will be based on the MIS 4 interval, during which we have no robust age
111
control within MD99-2198. We therefore present our discussion in terms of early, mid or late
112
MIS 4 etc. Fortunately, records obtained from individual cores can be compared together
113
unambiguously and several of our records have been obtained in such a way.
114
The L* reflectance index (a measure of sediment brightness) may be used to
115
distinguish sedimentological components such as free and bound Fe, CaCO3, Fe-minerals
116
(e.g. goethite), and clay [Rogerson et al., 2006]. A high-resolution record of core reflectance
117
(similar to L*) from the western tropical Atlantic was used to identify a link between
118
sediment reflectance changes in the Cariaco Basin (northern coastal Venezuela) to Greenland
119
ice core δ18O changes, thereby demonstrating a clear linkage of the tropical hydrological
120
cycle with high northern latitude climate [Peterson et al., 2000]. We note that the record of
121
L* from MD99-2198 [Hüls and Zahn, 2000] reveals a similar relationship with Greenland
122
temperature when placed on our timescale (Fig. 3).
5
123
Age control for ODP 1063 is explained in detail in a study by Thornalley et al. (2013)
124
and briefly was obtained by tuning core reflectance and magnetic susceptibility to records of
125
orbital precession and obliquity [Grützner et al., 2002], with further refinement by tuning the
126
planktic δ18O record to NGRIP δ18O (Fig. 4). Age control for RC24-01 was initially based on
127
a low-resolution δ18O record (derived from the thermocline dwelling N. dutertrei) from the
128
same core [Verardo and McIntyre, 1994], which was tuned to SPECMAP [Martinson et al.,
129
1987]. The age model was then fine-tuned using a new high resolution δ18O record from G.
130
ruber, and the consistency of the age model was checked by comparing the % CaCO3 record
131
of RC24-01 to that of ODP 1063 [Thornalley et al., 2013] (Fig. 4).
132
133
MD99-2198
134
Initially, porosity data from MD99-2198 (Labeyrie and Zahn, 2005) were used to identify
135
two anomalous intervals within the section of interest (Fig. S1). On the basis of very large
136
increases in porosity, these two intervals (1006-1024 cm and 1378-1430 cm) were interpreted
137
as representing episodes of core slumping (material within these intervals was not preserved
138
in the core barrel due to a lack of consolidation). We therefore modified the core depth scale
139
to exclude these intervals. All records are based on the modified core depth. The initial
140
selection of the depth interval of MD99-2198 which corresponded to MIS 5a-3 was made on
141
the basis of a long (~150 ka) record of benthic δ18O measured on the same core (R. Zahn,
142
unpublished data) (See Fig. S1). The age model was then fine-tuned by tuning a new high-
143
resolution record of planktic foraminiferal δ18O (Globigerinoides ruber, white, picked from
144
250-315 μm of the coarse fraction) to NGRIP δ18O (Main text Fig. 3).
145
6
146
RC24-01
147
Initial selection of the depth interval of RC24-01 which corresponded to MIS 5a-3 was made
148
on the basis of a long (~140 ka) planktic δ18O record (Neogloboquadrina dutertrei),
149
measured on the same core (Verardo and McIntyre, 1994) (Fig. S2). Verardo and McIntyre
150
(1994) established isotopic events in their δ18O record according to the critieria of Pisias et
151
al., (1984), and tuned their planktic δ18O record to the SPECMAP stacked oxygen isotope
152
record of Imbrie et al., (1984), using the chronology established by Martinson et al., (1987)
153
(Fig. S2, upper panel). High-resolution samples were then taken across the depth interval
154
approximately corresponding to MIS 5a-4. We then measured δ18O for these samples, on the
155
planktic foram Globigerinoides ruber (white), and used this record to fine-tune the age model
156
(Fig. S2, lower panel).
157
158
Samples for Neodymium isotope ratio measurements
159
Core MD99-2198 (12.09˚N; 61.23˚W; 1330m water depth) was sampled at 10cm intervals
160
between 1040 and 1110cm, based on the approximate position of the MIS 5/4 transition from
161
low-resolution benthic δ18O data (R. Zahn, unpublished data). Additional core samples were
162
taken at ~100cm intervals either side of the MIS 5a/4 transition. Two samples showed a low
163
ion beam strength (1-1.3 V), with significant isobaric interference from 144Sm. These samples
164
could not produce a viable value of εNd and were abandoned. One sample, identified with an
165
asterisk in paper Table 1, had an ion beam that was significantly smaller than for standards,
166
and hence a propagated error is reported to reflect this difference.
167
168
169
170
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Supplementary references
BACON, M. P. (1984), Glacial to interglacial changes in carbonate and clay sedimentation in the
Atlantic Ocean estimated from 230Th measurements, Chemical Geology, 46(2), 97-111.
BRADTMILLER, L. I., ANDERSON, R. F., FLEISHER, M. Q., and BURCKLE, L. H. (2007), Opal burial in the
equatorial Atlantic Ocean over the last 30 ka: Implications for glacial-interglacial changes in
the ocean silicon cycle, Paleoceanography, 22(4), PA4216, doi: 10.1029/2007PA001443.
CARDINAL, D., ALLEMAN, L. Y., DE JONG, J. ZIEGLER, K., and ANDRÉ, L. (2003), Isotopic composition
of silicon measured by multicollector plasma source mass spectrometry in dry plasma mode,
Journal of Analytical Atomic Spectrometry, 18(3), 213-218.
CHESTER, R., and HUGHES, M. (1967), A chemical technique for the separation of ferro-manganese
minerals, carbonate minerals and adsorbed trace elements from pelagic sediments Chemical
Geology, 2, 249-262.
FRANCOIS, R., FRANK, M. RUTGERS VAN DER LOEFF, M. M. and BACON, M. P. (2004), 230Th
normalization: An essential tool for interpreting sedimentary fluxes during the late
Quaternary, Paleoceanography, 19(1), PA1018, doi: 10.1029/2003pa000939.
GEORG, R. B., Reynolds, B. C., FRANK, M. and HALLIDAY, A. N. (2006), New sample preparation
techniques for the determination of Si isotopic composition using MC-ICPMS, Chemical
Geology, 235, 95-104.
HENDRY, K. R., LENG, M. J., ROBINSON, L. F., SLOANE, H. J., BLUSZTJAN, J. RICKABY, R. E. M., GEORG,
R. B. and HALLIDAY, A. N. (2011), Silicon isotopes in Antarctic sponges: an interlaboratory
comparison, Antarctic Science, 23(1), 34-42.
IMBRIE, J., HAYS, J. D., MARTINSON, D. G., MCINTYRE, A., MIX, A. C., MORLEY, J. J., PACES, N.G.,
PRELL, W. L. & SHACKLETON, N. J. (1984) The orbital theory of Pleistocene climate: Support
from a revised chronology of the marine δ18O record, in Milankovitch and Climate, Part I,
edited by A. Berger et al., pp. 269-305, D. Reidel, Norwell, Mass.
LABEYRIE, L. & ZAHN, R. (2005) Physical properties of sediment core MD99-2198., Pangaea.
MARTINSON, D. G., PISIAS, N. G., HAYS, J. D. , IMBRIE, J., MOORE Jr., T. C. and SHACKLETON, N. J.
(1987) Age dating and the orbital theory of the ice ages: Development of a high-resolution 0
to 300,000-year chronostratigraphy. Quaternary Research, 27(1), 1-29.
PISIAS, N. G. & LENIN, M. (1984) Milankovich forcing of the oceanic system: Evidence from the
northwest Pacific, in Milankovitch and Climate, Part I, edited by A. Berger et al., pp. 307-330,
D. Reidel, Norwell, Mass.
TANAKA, T., et al. (2000), JNdi-1: a neodymium isotopic reference in consistency with LaJolla
neodymium, Chemical Geology, 168(3–4), 279-281.
VERARDO, D. J., and MCINTYRE, A. (1994) Production and Destruction: Control of Biogenous
Sedimentation in the Tropical Atlantic 0-300,000 years B.P. Paleoceanography, 9(1), 63-86.
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Supplementary Data Tables
215
Table S1 - MD99-2198 Nd isotope data
Depth (cm)
Age (ka)
143Nd/144Nd
2  S.E.
εNd
2  S.D.
850-851
44.57
0.512116
0.000010
-10.2
0.2
950-951
54.93
0.512159
0.000012
-9.3
0.2
1024-1025
62.50
0.512121
0.000006
-10.1
0.3
1034-1035
63.52
0.512124
0.000010
-10.0
0.3
1044-1045
64.54
0.512174
0.000014
-9.1
0.3
1054-1055
65.56
0.512130
0.000010
-9.9
0.3
1064-1065
66.58
0.512155
0.000008
-9.4
0.3
1084-1085
68.33
0.512127
0.000012
-10.0
0.3
1094-1095
69.42
0.512072
0.000012
-11.1
0.3
1114-1115
71.05
0.512087
0.000020
-10.8
0.3
1139-1140
72.64
0.512097
0.000018
-10.6
0.3
1249-1250
79.08
0.512088
0.000012
-10.7
0.3
1394-1395 *
87.04
0.512103
0.000014
-10.4
0.6
216
217
All reported 143Nd/144Nd ratios have been normalised to the recommended JNdi value
143
Nd/144Nd
218
of Tanaka et al. (2000). Epsilon Nd values denote the deviation of measured
219
values from the bulk Earth value (CHUR=0.512638) in parts per 10,000. The external
220
reproducibility (2σSD) is reported based on repeat JNdi analyses of the day. For one sample
221
(*), the ion beam was significantly smaller than for standards and hence a propagated error is
222
reported to reflect this difference.
9
223
Table S2 - MD99-2918 Si isotope data. External reproducibility for δ30Si values is +/- 0.23
224
per mil.
Mean δ30Si
~[Si(OH)4]
Depth (cm)
Age (ka)
(all analyses)
(μM)
850-851
44.57
-0.86
11.6
950-951
54.93
-1.13
16.0
1024-1025
62.50
-0.74
9.8
1034-1035
63.52
-0.73
9.6
1044-1045
64.54
-0.96
13.1
1054-1055
65.56
-0.43
5.5
1064-1065
66.58
-0.39
5.0
1074-1075
67.60
-0.35
4.6
1084-1085
68.33
-0.23
3.1
1094-1095
69.64
-0.12
1.9
1114-1115
71.05
-0.12
1.8
1139-1140
72.64
-0.57
7.4
1184-1185
75.51
-0.58
7.5
1249-1250
79.09
0.28
-2.2
1394-1395
87.03
0.15
-1.0
225
226
227
228
10
229
Table S3 - RC24-01 radionuclide data
Depth
Age (ka)
% opal
230Th-norm
error +/-
F(g/cm2/ka)
(cm)
Uauth
error +/-
(dpm/g)
284.5
60.92
6.65
1.259
0.023
0.880
0.012
293.5
62.88
8.40
1.220
0.029
1.257
0.017
302.5
64.84
7.78
1.270
0.035
1.781
0.026
310.5
66.58
10.59
1.062
0.038
2.644
0.034
318.5
68.32
10.38
1.205
0.040
2.818
0.039
325.5
69.84
8.94
1.239
0.047
2.046
0.027
333.5
71.59
7.09
1.321
0.050
1.486
0.020
344.5
74.97
4.62
1.332
0.053
1.298
0.018
351.5
78.04
5.32
1.299
0.045
1.097
0.017
353.5
78.92
4.95
1.331
0.037
0.778
0.013
365.5
84.18
3.67
1.403
0.043
0.304
0.005
373.5
87.16
3.73
1.538
0.041
0.133
0.002
230
231
232
233
234
235
236
11
237
238
Supplementary Figures
Figure S1
239
12
240
Figure S2
241
242
243
244
245
13
246
Figure S3
247
248
249
250
251
252
253
254
255
256
14
257
Supplementary Figure annotations
258
259
Figure S1. Top panel, blue curve: Plot of porosity (% vol) (Labeyrie and Zahn, 2005) against
260
original MD99-2198 core depth. Intervals marked in the top panel with shaded bands were
261
interpreted as representing core slumps, and were consequently excluded from the record,
262
with core depth adjusted accordingly. All records are based on the modified core depth, and
263
the planktic δ18O record of MD99-2198 (this study, green curve) is plotted against the
264
modified record of benthic δ18O (red curve) from Globigerinoides sacculifer (R. Zahn,
265
unpublished data).
266
267
Figure S2. (a) NGRIP δ18O record (grey curve); overlain by RC24-01 planktic δ18O record
268
from N. dutertrei (Verardo and McIntyre, 1994) (red curve), with offset scale (in red) to
269
reflect differences in δ18O between N. dutertrei and G. ruber; and RC24-01 planktic δ18O
270
record from G. ruber (this study) (blue curve). Both sets of RC24-01 planktic δ18O data in (a)
271
are plotted on the chronology of Verardo and McIntyre (1994), tie points are shown in
272
magenta.
273
(b) NGRIP δ18O record (grey curve); overlain by RC24-01 planktic δ18O record from N.
274
dutertrei (Verardo and McIntyre, 1994) (red curve), with offset scale (in red) to reflect
275
differences in δ18O between N. dutertrei and G. ruber; and RC24-01 planktic δ18O record
276
from G. ruber (this study) (blue curve). Both sets of RC24-01 planktic δ18O data in (b) are
277
plotted on the revised chronology from this study, old tie points (the same as Verardo and
278
McIntyre, 1994) are shown in magenta; new tie points based on the revised chronology are
279
shown in orange.
15
280
(c) sedimentation rates (cm/ka) derived from the chronology established by this study, based
281
on fine tuning the planktic δ18O record from G. ruber to NGRIP (solid line), compared on the
282
same y-axis to the sedimentation rates (cm/ka) derived from the chronology established by
283
Verardo and McIntyre (1994) (dashed line).
284
285
Figure S3. A plot of the range of expected silicic acid concentrations (orange shaded area),
286
given the predicted δ30Si variability of +1.5 – +2 ‰ in AAIW [Hendry et al., 2010; de Souza
287
et al., 2012]. Calculated values of paleo-[Si(OH)4] assuming a δ30Si for AAIW of 1.5 ‰ [de
288
Souza et al., 2012], are shown in black.
289
290
291
292
293
294
295
296
16