ARTICLE IN PRESS
Quaternary Science Reviews ] (]]]]) ]]]–]]]
Extended orbitally forced palaeoclimatic records from the equatorial
Atlantic Ceara Rise
Heiko Pälikea,, Julia Frazierb, James C. Zachosb
a
National Oceanography Centre, Southampton, School of Ocean and Earth Science, European Way, Southampton SO14 3ZH, UK
b
Earth Sciences Department, University of California, Santa Cruz, CA 95064, USA
Accepted 20 February 2006
Abstract
We extend existing high-resolution Oligocene–Miocene proxy records from Ocean Drilling Program (ODP) Leg 154. The extended
record spans the time interval from 17:86 to 26.5 Ma. The data are age calibrated against a new astronomical solution that affords a reevaluation of the intricate interaction between orbital (‘‘Milankovitch’’) forcing of the climate and ocean system, and the fidelity with
which this forcing is recorded in oxygen and carbon stable isotope measurements from benthic foraminifera, and associated lithological
proxy records of magnetic susceptibility, colour reflectance, and the measured sand fraction. Our records show a very strong continual
imprint of the Earth’s obliquity cycle, modulate in amplitude every 41 ka, a very strong eccentricity signal in the carbon isotope records,
and a strong, but probably local, imprint of climatic precession on the coarse fraction and magnetic susceptibility records. Our data
allowed us to evaluate how the interaction of long, multi-million year beats in the Earth’s eccentricity and obliquity are implicated in the
waxing and waning of ice-sheets, presumably on Antarctica. Our refined age model confirms the revised age of the Oligocene–Miocene
boundary, previously established by analysis of the lithological data, and allows a strong correlation with the geomagnetic time scale by
comparison with data from ODP Site 1090, Southern Ocean.
r 2006 Elsevier Ltd. All rights reserved.
1. Introduction
Ocean Drilling Program (ODP) Leg 154 data (Curry
et al., 1995; Shackleton et al., 1997) so far provide the best
developed set of demonstrably astronomically forced
lithological and climatic proxy data cycles spanning the
Neogene, and extending as far back as the late Eocene
(Shackleton et al., 1999). A main strength of palaeo-climate
series derived from ODP Leg 154 sediments is that
astronomically calibrated age models have been developed
based on lithological (Shackleton and Crowhurst, 1997;
Weedon et al., 1997; Shackleton et al., 1999) and
geochemical (Flower et al., 1997; Pearson et al., 1997)
variations, that allow independent geochemical proxies to
be placed on a time scale of high accuracy, potentially at
the climatic precession period ð21 kaÞ. Multiple studies
have exploited the data sets generated from ODP Leg 154
for palaeoceanographic (Shackleton and Crowhurst, 1997;
Corresponding author. Tel.: +44 2380 593638; fax: +44 2380 593052.
E-mail address: H.Palike@soton.ac.uk (H. Pälike).
0277-3791/$ - see front matter r 2006 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2006.02.011
Zachos et al., 1997, 2001a,b; Paul et al., 2000; Flower et al.,
2006), biostratigraphic (Raffi, 1999; Shackleton et al.,
2000) and cyclostratigraphic (Shackleton et al., 1999;
Pälike and Shackleton, 2000; Billups et al., 2004; Pälike
et al., 2004; Lourens et al., 2005) purposes, and are now
being exploited for the interpretation of data from more
recent drilling expeditions (Flower and Chisholm, 2006).
ODP Leg 154 sediments offered the opportunity to study
palaeoceanographic processes at a resolution and fidelity
previously limited to the Plio- and Pleistocene, over several
million years. Some of these studies investigated the close
relationship between the cyclostratigraphic record and
macroscopic (multi-million year) features of astronomical
calculations, made possible by the extended calculations of
Laskar et al. (1993). These astronomical calculations have
recently been evaluated (Néron de Surgy and Laskar, 1997;
Laskar, 1999) and improved (Varadi et al., 2003; Laskar
et al., 2004). The age models from ODP Leg 154 were
adjusted to these new calculations (Pälike et al., 2004),
demonstrating a better fit with these newer solutions
compared to the originally used one of Laskar et al.
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2
(1993), and make it opportune to re-visit the impact of this
re-tuning on palaeoclimatic proxy series, as well as
formally establishing the basis for the re-calibration by
publishing the full data series (see supplementary material).
In addition, here we extend the original stable isotope
record that was established across the Oligocene/Miocene
boundary (Paul et al., 2000; Zachos et al., 2001b) both upand down-core, and focus in more depth on the implications of the associated percent sand fraction ð465 mmÞ
record, with a series of significant results.
2. Methods
The data sets presented here will be made available in
electronic form at a designated data repository.1 The
generation of new benthic stable isotope data ðd18 O; d13 CÞ
and percent sand fraction ð465 mmÞ follows the methods
described by Paul et al. (2000), and were also measured in
Santa Cruz. The initial study by Paul et al. (2000)
generated data from two holes: 926B (3 430 N, 42 540 W,
3.6 km water depth) and 929A (5 580 N, 43 440 W, 4.36 km
water depth), which were used to generate a spliced record
from these two Holes, covering core gaps and identified
local slumped horizons. Here, we focused on extending
upward and downward the record from ODP Site 926B,
the shallower of the two localities, and the one with the
higher sedimentation rates. The new record extending
downward is based on analyses of a mixture of Cibicidoides species including C. mundulus and C. cresbi, while
the new record extending upward is based entirely on
analyses of C. mundulus. The complete data we present here
were sampled from Core 926B-35X downwards to Core
926B-59X, excluding intervals that were identified as local
slumps (Shackleton et al., 1999), and sampled every 10 cm.
The data set also comprised previously generated data
from a sub-interval (Paul et al., 2000; Zachos et al., 2001b).
Sedimentation rates for this interval were previously
estimated by Weedon et al. (1997) and astronomically
calibrated by Shackleton et al. (1999): they range between
25 and 35 m/Ma in the younger time interval, and
between 20 and 30 m/Ma in the older time interval of
the record, resulting in a median sample step of 3:5 ka.
On our new time scale, this entire record now spans the
interval 17:86226:5 Ma (17:77226:24 Ma on the time
scales of Shackleton et al., 1999, 2000). The ODP 926B
record contains short core breaks, and the data are linked
into a composite section by comparison with a spliced
combined record of susceptibility and reflectance obtained
from several ODP 154 Sites (Shackleton et al., 1999),
normalised to zero mean and unit standard deviation as
described and used by Pälike et al. (2004). The extended
data set used in this study is presented in Fig. 1 against
Hole 926B core depth in metres composite depth (mcd)
(Curry et al., 1995; Shackleton et al., 1999). The stacked
susceptibility and colour reflectance record was obtained
1
http://www.ncdc.noaa.gov/oa/ncdc.html.
from multiple Sites, but re-adjusted to the 926B depth
scale. It was generated such that positive excursions in this
record coincide with positive excursions in the susceptibility record obtained directly from Hole 926B. The coarse
fraction data for ODP Hole 926B shown in Fig. 1
supplement those published previously from ODP Site
929 (Zachos et al., 1997; Paul et al., 2000), which had a
deeper palaeo-depth. Prior to further analysis, we prepared
the percent p
coarse
fraction record byffi using the transformaffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
tion arcsinð %coarse fraction=100Þ. This transformation
was applied before further time series analysis in order to
achieve more normally distributed values from data that
are clipped at zero percent (Weedon, 2003, p. 54).
2.1. Tuning strategies
Significant previous work was undertaken to combine
geological data from ODP Leg 154 with calculated
variations of the Earth’s orbit in order to achieve an
orbitally calibrated age model. This work comprised
several important steps. First, sequences from different
ODP Leg 154 sites and holes were cross-correlated in order
to obtain a continuous record from individual cores, and in
order to evaluate the continuity of this ‘‘spliced’’ data set.
The correlation was accomplished on the basis of highresolution lithological data (chiefly, magnetic susceptibility, colour reflectance, down-hole logging data, as well as
high-resolution biostratigraphic data). This extensive work,
and an evaluation of the stratigraphic continuity of the
combined records, was achieved by Weedon et al. (1997)
and Shackleton et al. (1999). From these composite spliced
records, the next step consisted of analysing the character
of any possible astronomical cyclicity on the basis of
existing bio-stratigraphic age assignments. Shackleton et al.
(1999) used the observation of a dominating obliquity
signal and a weaker but discernible climatic precession
signal to generate an appropriate astronomical target
curve, and used the longer-term amplitude modulation of
obliquity (1:2 million year long amplitude envelope) and
climatic precession (400 ka envelope) to evaluate the
fidelity of the generated age model within ‘‘hierarchies’’ of
orbital cycle time scales.
The resulting age model, generated on the basis of
lithological variations, allowed other proxy data to be
placed on this time scale, permitting the calculation of
sedimentary fluxes and evaluation of rates of evolutionary
change with a resolution and accuracy that is far better
than can be achieved by traditional geological means.
Shackleton et al. (1999) found that the 1.2 million year long
amplitude modulation cycle of obliquity indicated that the
age of the Oligocene/Miocene boundary be moved from its
previous age of 23.8 Ma (Cande and Kent, 1995) to
22.9 Ma (Shackleton et al., 2000), and that the amplitude
variation of climatic precession would place a further
constraint on any alternative age model (allowing shifting
by 405 ka multiples). A different view for the correct age
of the Oligocene/Miocene boundary was proposed by
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(a)
(b)
(c)
SusRef
0
(e)
-2
-6
(SI x 10 )
15
10
(f)
13
δ C
SUS
(normalised)
2
(d)
3
5
(‰)
2
0
1
0
-1
350
450
Mi-1
Hole 926B Depth (mcd)
400
500
550
20
40
60
Lightness
(%)
0
10
20
30
Coarse fraction
(%)
2
1
0
18
δ O
(‰)
Fig. 1. ODP Leg 154 data used in this study plotted versus ODP 926B core depth. The ODP 926B record contains short core breaks, and the data are
linked into a composite section by comparison with a spliced combined record (a) of susceptibility and reflectance (‘‘SusRef’’), normalised to zero mean
and unit standard deviation as described and used by (Pälike et al., 2004). Also plotted are (b) the colour reflectance parameter L* (lightness) and (c)
magnetic susceptibility (‘‘SUS’’), calculated by gaussian interpolation at stable isotope sample points. Previously published benthic foraminifer data from
ODP Site 926B (Paul et al., 2000; Zachos et al., 2001b) were extended and are plotted with (d) the 463 mm coarse fraction data. Benthic d13 C (e) and d18 O
(f) are plotted referenced against VPDB. Mi–1 denotes the position of a prominent glacial event previously discussed (Miller et al., 1991; Zachos et al.,
2001b).
Wilson et al. (2002). However, the findings of this study
have been questioned (Channell and Martin, 2003), and
additional studies confirmed the new correlation of
Shackleton et al. (2000) for sediments from other localities
(Billups et al., 2004; Pälike et al., 2004), subject to readjustment to the new astronomical solution of Laskar
et al. (2004).
Importantly, the time scale from Shackleton et al. (1999)
was developed using an extension of the calculations that
were done by Laskar et al. (1993), which were found to
have an inherent limit for the duration over which they can
be considered valid into the past (Laskar, 1999). Very
recently, an updated and much improved calculation of
Earth’s orbital variations was published (Laskar et al.,
2004). Pälike et al. (2004) demonstrated using the
lithological data from ODP Legs 154 and 199 that this
new calculation fits better with the geological data. Here,
we report the results of re-adjusting the astronomical age
calibration of ODP Leg 154 data to the new astronomical
solution, which moves the age of Oligocene/Miocene
boundary back in time by 100 ka, to a new age of
23.03 Ma (Lourens et al., 2005), chiefly controlled by the
difference in 100 ka eccentricity cycle amplitudes.
The re-tuning of magnetic susceptibility data to the new
Laskar et al. (2004) solution was done by aligning data and
target at the climatic precession frequency, constrained by
the 100 ka amplitude modulation of precession by
eccentricity. This assumption of a zero-phase difference
at the precession frequency between astronomical solution
and geological data results in a phase lag at the obliquity
frequency of 6–7 ka during the late Oligocene (Billups
et al., 2004). We note here that the new astronomical
solution of Laskar et al. (2004) uses current day values for
tidal dissipation and dynamical ellipticity (Pälike and
Shackleton, 2000), the established phase relationships for
obliquity and climatic precession would be altered when
different histories for the tidal dissipation and dynamical
ellipticity terms were used. A subsequent iteration applied
this phase lag at the obliquity frequency for the calculated
ETP curve, generating the overall best-fitting target curve.
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4
The zero-phase assumption at precession stems from the
evidence that the strong precession signal arises from local
climatic processes on the adjacent continent which
modulate the terrigenous input, and possibly local productivity and the preserved sediment coarse fraction. The
lag at the obliquity frequency presumably arises from the
slow response of the varying Antarctic continental ice
sheet, and for d13 C on the long residence time of carbon
and nutrients in the ocean. The phase relationship used
follows that of Shackleton and Crowhurst (1997). They
interpreted magnetic susceptibility values as a proxy for the
terrigenous supply and correlated maxima in susceptibility
with maxima in Southern Hemisphere summer insolation
(minima in Northern Hemisphere insolation), based on the
observations that otherwise positive d18 O values would
correspond to Northern Hemisphere insolation minima,
and the obliquity response would be out of phase. In this
scenario, increased terrigenous input through the Amazon
system would be driven by maxima in Southern Hemisphere insolation. They also argued, though, that in the
(h)
(g)
(f)
ETP
-3
17
18
0
Oligocene the uncertainties in the orbital calculations are
still at an order (Pälike and Shackleton, 2000; Lourens
et al., 2000) where age errors that arise from such phase
assumptions are smaller than the error that arises from our
incomplete knowledge of the dissipative effects (tidal
dissipation) in the orbital solutions. The data sets are
plotted against our new time scale in Fig. 2, together with
the astronomical templates from Laskar et al. (2004), as
well as a comparison of the previous time scale of
Shackleton et al. (1999) that was generated with the
astronomical solution by Laskar et al. (1993) and applied
by Zachos et al. (2001b) to a sub-set of the stable isotope
data. ‘‘ETP’’ refers to an artificial mix of eccentricity,
obliquity (tilt) and climatic precession, that are combined
by normalising these three components (subtracting their
mean and dividing by their standard deviation), and then
adding these with a set of weightings determined by the
geological data. In this case the eccentricity component was
accentuated as the stable isotope data have a large
eccentricity imprint. We note that ETP resembles Northern
(e)
(d)
(c)
(b)
13
SusRef
δ C
(normalised)
(‰)
3
2
0
-2
2
1
(a)
18
δ O(‰)
(Zachos et al. 2001b)
0
2
-1
1
0
46
MiC5En
19
49
MiC6n
20
Age (Ma)
52
MiC6An
21
55
MiC6Bn
22
Mi-1
23
58
OlC6Cn
24
61
OlC7n
25
54
Ol-
26
C8n
-1
0
1
400 kyr ecc
22
24
Obliquity (deg)
0.0
0.2
0.4
0.6
Coarse fraction
-1
sin (%/100)
0.5
2
1
0
18
δ O
(‰)
Fig. 2. ODP Leg 154 data plotted versus revised astronomical age (Laskar et al., 2004). Shown are (a) the previously published spliced d18 O dataset from
Sites 926 and 929, age calibrated using a previous astronomical solution (Zachos et al., 2001b; Laskar et al.,p1993),
against
new extended d18 O (b) and d12 C
ffiffiffiffiffiffiffiffiffiffiffiffiffiffi
ffi
(c) data, using a revised astronomical solution and time scale (Laskar et al., 2004). (d) Shows an arcsinð %=100Þ transformation of the age-calibrated
coarse fraction percentage data. This transformation was applied before further time series analysis in order to achieve more normally distributed values
from data that are clipped at zero percent (Weedon, 2003, p. 54). (e) Shows age-calibrated combined susceptibility and lightness data. (f) Shows Earth’s
obliquity and 1:2 Ma obliquity amplitude modulation from Laskar et al. (2004). (g) Calculated mix of eccentricity, obliquity (tilt), and climatic precession
(‘‘ETP’’), using Laskar et al. (2004), with relative amplitudes reflecting those of the data, and 6:4 ka lagged obliquity component (Billups et al., 2004). (h)
400 ka bandpass filters (0:5 Ma1 gaussian bandwidth) applied on ETP (black), d18 O (blue) and d13 C (green). Red tick marks and codes mark the
position, count and identification of the absolute number of 400 ka cycles as first proposed by Wade and Pälike (2004). Thin cross-lines mark the position
of obliquity amplitude minima, thick green cross-lines mark the position of long-term eccentricity amplitude minima.
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Hemisphere summer insolation maxima at climatic precession minima, depending on the definition of the climatic
precession phase angle (Berger et al., 1993). In Fig. 2(h) we
have also marked the position of absolute 405 ka cycle
codes, as first described by Wade and Pälike (2004),
supplementing the code with a letter abbreviation for the
geological epoch (Mi for Miocene, Ol for Oligocene) and
an indication within which magnetic reversal interval this
cycle falls for existing astronomically calibrated sections
(Lourens et al., 2005), following the convention of Cande
and Kent (1995).
2.2. Spectral analysis
We conducted different spectral analysis methods on the
astronomical curves, the benthic stable isotope data, the
coarse fraction record, and the spliced and stacked
susceptibility and reflectance (‘‘SusRef’’) record. The
interval analysed extends from 17.86 to 26.5 Ma. Prior to
analysis, the data were re-sampled at the median sample
resolution every 4 ka. Data sets were linearly detrended,
and long-term trends (periods longer than 1 Ma) were
removed with a gaussian notch filter, using version 2 of the
software package AnalySeries (Paillard et al., 1996). Prior
to analysis, we multiplied the two stable isotope series with
1, consistent with Zachos et al. (2001b), and aligning
oxygen isotope maxima with temperature minima (or icevolume maxima). We used three main tools for spectral
analysis. Power spectra were calculated using the Multitaper method with a timestep bandwidth product of 4 and
6 tapers, using the SSA-MTM Toolkit v4.2 (Ghil et al.,
2002). This software was also used to fit red-noise
significance limits at the 90% and 99% levels. We used
the Blackman–Tukey method in AnalySeries to calculate
cross-spectral coherence and phase relationships between
an artifical mix of eccentricity, obliquity (tilt), and climatic
precession (ETP). For the coherency, we determined the
95% confidence interval testing the non-zero hypothesis.
We also computed evolutive Multitaper-method spectra for
the data, using a sliding 1.6 Ma window stepped at 50 ka
intervals. For the cross-spectral analysis the resulting
frequency bandwidth is 0:86 Ma1 .
3. Results
3.1. ODP Leg 154 data re-calibrated to new astronomical
calculation
Our first result concerns the re-adjustment of the ODP
Leg 154 data set to the refined astronomical solution of
Laskar et al. (2004). Fig. 2 (a) and (b) shows a comparison
of the benthic oxygen isotope data as used by Zachos et al.
(2001b), using the time scale of Shackleton et al. (1999), as
well as the re-adjusted data on the time scale of our study.
In terms of the Shackleton et al. (1999) time scale, our new
ages are approximately 20 ka older in the interval prior to
19:71 Ma (using Shackleton et al., 1999 ages). This age
5
shift towards slightly older ages then gradually increases
between 19:71 and 20.15 Ma and reaches an average
offset of approximately 90 ka for the remainder of the data,
except one interval between 23:35 and 24.21 Ma, where
this reaches 160 ka. This difference is chiefly controlled
by the different nature of short eccentricity cycles in the
two different target curves used (for the offset of 90 ka or
one eccentricity cycle), and some additional small refinements in the original splice that were done after the
publication of Shackleton et al. (1999) (N. J. Shackleton,
pers. comm. 2002). We note that on our new re-adjusted
time scale, evaluated in detail by Pälike et al. (2004), the
position of the oxygen and carbon isotope excursions
previously designated as ‘‘Mi–1’’ again fall into a ‘‘node’’
of minimum obliquity amplitude variation, and coincide
approximately with the position of the Oligocene/Miocene
boundary (Shackleton et al., 2000). The revised ages
presented in our study were the basis for the age given
for the O/M boundary in Lourens et al. (2005). Some of the
additional magnetic reversal ages given in Lourens et al.
(2005) were derived from Australia–Antarctic spreading
distances, which are said to improve the original approach
of fitting a smooth spline curve to a smaller number of tiepoints as was done by Cande and Kent (1995). The fact
that Mi–1, here designated as long eccentricity cycle 58,
falls into an obliquity amplitude node, is consistent with
the original placement in Zachos et al. (2001b). However,
the overall influence of the orbital geometry on glaciations
is more complex than previously argued, and will be
discussed below.
3.2. Time evolution of spectral characteristics
The combination of a refined astronomical age scale
together with an extension of the time coverage of the data
allows a re-evaluation of the spectral characteristics
contained in the data sets, and an enhanced analysis of
the relationship between long-term astronomical forcing
and the encoding of this forcing in the different data sets.
The spectral characteristics of the data and the astronomical solution of Laskar et al. (2004) are summarised in
Fig. 3. For all data sets, the 40 ka obliquity cycle is the
one that most significantly stands out above the background spectrum. This is the reason why the age model of
Shackleton et al. (1999) used this cycle as the pre-dominant
tuning target, in combination with the climatic precession
imprint in the SusRef data set. Both stable isotope time
series contain significant short and long eccentricity cycles
(110 and 405 ka). Interestingly, of the palaeoceanographic proxy time series, the coarse fraction record
contains significant spectral peaks at the three climatic
precession frequencies, which were also established in the
spliced record of Zachos et al. (2001b). We can now reevaluate the fidelity of the 1:2 Ma long obliquity
amplitude modulation cycle that was already established
for the SusRef record (Shackleton et al., 1999; Pälike et al.,
2004). Throughout the record, the oxygen isotope data
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Power
1
0.1 0.01 0.9 0.4
19
20
21
22
24
25
ODP 926B δ C
13
Bandwidth
Confidence
50
99%
90%
19
20
21
22
Coherency
23
24
25
ODP 926B asin coarse%/100
19
Confidence
99% 22
90% 23
“1.2 Myr” obliquity
41
20
(54)
10
96
126
405 (d)
0
60
ODP 926B δ18O
Bandwidth
Confidence 19
99%
90% 22
23
ODP 926B SusRef
40
(29)
30
41
(54)
20
10
(b)
96
126
405 (e)
0
60
50
Frequency (Myr -1)
1
Bandwidth
40
30
Power
0.4 0.9 0.01 0.1
“2.4 Myr” eccentricity
Bandwidth
Confidence
50
99%
90%
Frequency (Myr -1)
23
Period (kyr)
Frequency (Myr -1)
60
(a)
Age (Ma)
Age (Ma)
Coherency
Period (kyr)
6
−δ13C −δ18 O
Astronomy (La2004)
Period (kyr)
19
22
23
40
(29)
“1.2 Myr” obliquity
30
41
20
(54)
10
96
126
0
405
50 40 30 20 10 0 -10 19
−δ13C
-δ18O
-SusRef
coarse(%)
“2.4 Myr” eccentricity
(c)
Lag (kyr)
(f)
20
21
22
23
24
25
Age (Ma)
180°
lag
90°
0°
-90°
Phase (°) relative to ETP
-180°
lead
Fig. 3. (a) and (b) log-power spectrum of benthic d13 C and d18 O time series, determined by the multi-taper method with robust red-noise background
for significance testing (Ghil et al., 2002). Prior to analysis the record was linearly detrended, and a Gaussian notch filter removed periods longer than
1 Ma. The time series was interpolated at equal time steps, using the average sample resolution ð4 kaÞ. Also shown are coherency estimates between
data and astronomical eccentricity, tilt and precession mix (ETP), using the Blackman–Tukey method, with 437 lags (20% of series from 17.86 to 26.5 Ma).
Confidence levels are shown for the spectra (solid 99%, dashed 90%), and the coherency estimates (95%). All Milankovitch frequencies are represented
and significant. Oxygen isotope data show a relatively stronger signal at the climatic precession frequencies in the younger half of the record. Also shown
are evolutive multi-taper method log-power spectra for d13 C, d18 O on the same frequency scale. (c) Blackman–Tukey cross-spectral phase estimates
between stable isotopes and astronomical ETP curve, converted from phase angles to lag times in ka, and shown with 95% confidence intervals (Paillard et
al., 1996). Both isotope series were inverted prior to analysis, following convention (Zachos et al., 2001b). The carbon and oxygen isotope data are exactly
out of phase at the obliquity frequency only. For precession the phase angles are essentially zero. For long eccentricity (405 ka) periods, there is a relative
lag of d13 C compared to d18 O of 25 ka, similar to previous results from the late Oligocene, including an apparently increasing lag of d13 C compared
to d18 O going from the shorter (96 ka) to the long (405 ka) eccentricity period (Zachos et al., 2001b; Billups et al., 2004). Also shown is the evolutive
spectrum of ETP, re-sampled at the resolution of the data. (d) and (e) Evolutive multi-taper method log-power spectra for transformed percent coarse
fraction and combined magnetic susceptibility and lightness data (Pälike et al., 2004), as in (a) and (b). (f) Cross-phase estimates for all data sets as phase
angles (within their 95% confidence envelopes).
display obliquity cycle amplitude variations co-varying
with the astronomical target curve. Minima in the
amplitude of the 41 ka cycle of the astronomical data
are indicated by dark shaded lines in Fig. 3. This pattern
also fits for the benthic carbon isotope record. The
evolutive spectrum of the coarse fraction record also
displays the 2:4 Ma long amplitude modulation of
eccentricity at the climatic precession frequency bands.
This feature is also apparent in the amplitude variation of
the 405 ka long eccentricity cycle for the benthic carbon
isotope data. Phase relationships between astronomical
and geological data are similar to previous results (Zachos
et al., 2001b), as shown in Fig. 3(c) and (f). The climatic
precession component of the tuning data set SusRef is in
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(a)
(d)
(e)
18
δ O(‰+0.64)
13
δ C (‰)
ODP 177 PMag
2
17
20
21
C
2
1
n
.1
7n
C 2n
24
3
-1
n
1n
1n
n.
an ar. n.1
6C
6A 6A n 6B n
C 2n 3n
C 2 C 2
22
0
1
(Zachos et al. 2001a,b)
1n
n.
6A n
2
19
C
Age (Ma)
18
6n
C
We can now compare our extended tuned records with
one that was recently developed for a benthic stable isotope
data set from ODP Leg 177, Site 1090, Southern Ocean,
Atlantic sector (Billups et al., 2002, 2004), which also used
the new astronomical solution of Laskar et al. (2004) as a
tuning target. The comparison of our data set with that
from the Southern Ocean is shown in Fig. 4. The record
from ODP Leg 177 contained a well-defined data set of
magnetic polarity measurements (Channell et al., 2003),
and the ages obtained for ODP Leg 154 can be used to
refine the geomagnetic polarity time scale (Cande and
Kent, 1995). The age models for our present study and that
of Billups et al. (2004) were obtained independently, yet
show a close match, particularly of the carbon isotope
record (Fig. 4). As discussed by Billups et al. (2004), across
the Oligocene/Miocene interval the basin-to-basin offset of
carbon isotope values is virtually absent, even 5 million
years into the Miocene. The Southern Ocean oxygen
isotope values show a consistent heavy offset of 0:5%
with respect to the equatorial Atlantic values. Fig. 4 also
shows a comparison of our extended data set with the
global multi-site compilation of Zachos et al. (2001a). The
original data set from Zachos et al. (2001b) made up a large
portion of the global compilation curve back to 25 Ma on
our new time scale. As shown by Lear et al. (2004) using a
new data set from the equatorial Pacific, the apparent ‘‘late
Oligocene warming’’ that is suggested by the global
compilation is largely artefact of using ODP Leg 154 data
for ages younger than 25 Ma, and sub-Antarctic data for
the older part of the Oligocene. The stable isotope data
presented here show a much more gradual and smaller
decrease towards lighter values over more than 4 million
years before the onset of Mi–1. We provide a close-up of
the isotope data sets in Fig. 5, which compares the
(c)
ODP 154,177
23
3.3. Comparison with Southern Ocean record and global
compilation
(b)
n
n r.1
n
5E
5D 5D
C
C C
phase with the astronomical target. The obliquity is in
phase between the lithological data set and the coarse
fraction record with respect to lagged obliquity. The
benthic oxygen isotope data show a further lag of about
3–4 ka at the obliquity frequency, while carbon isotope
data are almost 1801out of phase with the oxygen isotope
data. This is in contrast to the eccentricity periods, where
oxygen and carbon isotopes co-vary, and also to observations from other localities spanning the same time interval
(Billups et al., 2004). In terms of absolute phase lags (in
ka), it is interesting to note that the lag of the carbon
isotopes in particular increases with increasing cycle
periods, an observation that is compatible with the
predictions of box models of the carbon cycle if forced
by astronomy, specifically with respect to the size of the
carbon reservoir and its preferential filtering of longer
periods due to the residence time of carbon in the oceans
(Pälike, unpublished results, using a modified version of the
Walker and Kasting, 1992 model).
7
25
26
2
1
13
0
δ C(‰)
-1
3
2
1
ODP 154, 177
18
(Zachos et al. 2001a,b) δ O (‰+0.64)
Fig. 4. Comparison of equatorial Atlantic ODP 154 data set with previous
global multi-site compilations (that include ODP 154 data) and previously
published data from the Southern Ocean: (a) magnetic reversal data from
ODP Site 1090 (Billups et al., 2004). (b) Previous d13 C data from ODP 154
as published in Zachos et al. (2001b), compared with a multi-site
compilation (Zachos et al., 2001a) that was age adjusted to new reversal
ages. The data from Zachos et al. (2001b) use an older astronomical
solution (Laskar et al., 1993; Shackleton et al., 1999). (c) d13 C data from
ODP 1090 (Billups et al., 2004) and independently age calibrated data
from this study, using the more recent astronomical solution Laskar et al.
(2004). (d) As in (c) but for d18 O. (e) As in (b) but for d18 O.
independently age-calibrated data sets from ODP Site
1090 (Billups et al., 2004), ODP Hole 929A (Zachos et al.,
2001b) and our study that uses data from ODP Hole 926B.
In most cases the oxygen and carbon isotope data show
that the three records agree down to the obliquity scale
level of variations. The new astronomical solution La2004
(Laskar et al., 2004) does still exhibit a node in the
obliquity amplitude variation, but this is not as pronounced during the Mi–1 event as previous astronomical
solutions. Thus, in order to explain the magnitude of Mi–1
compared to previous and following obliquity nodes a
further pre-conditioning, for example through a decline of
atmospheric pCO2 concentrations at this time is required
(Deconto and Pollard, 2003).
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24
(b)
1
(f)
2
2
Hole 926B
S. disbelemnos
Sphenolithus delphix
S. ciperoensis
δ C (‰)
0
Hole 929A
%Sphenoliths
0
20
3
3
1
0
1
23
Age (Ma)
2
2n
ODP 1090D,E
ODP 929A
ODP 926B
C7n.1n
3n
2n
2
C6Cn.1n
13
(e)
ODP 1090D,E
ODP 926B
1
ODP 177
PMag
(d)
ODP 154
Holes 926B,929A
(c)
ODP 154, Hole 926B
22
ODP 177, Holes 1090D,E
13
δ C (‰)
18
δ O (‰+0.64) Obliquity (deg)
(a)
ODP 177, Holes 1090D,E
18
δ O (‰+0.64)
8
6
4
2
0
Mi-1 “Event”
Eccentricity (%)
H. Pälike et al. / Quaternary Science Reviews ] (]]]]) ]]]–]]]
8
24
Miocene Oligocene
Fig. 5. Close-up of data around Oligocene–Miocene boundary data between 22.4 and 243 Ma on our revised time scale. (a) Eccentricity (b) obliquity
calculations from Laskar et al. (2004), including the 1:2 Ma obliquity amplitude envelope. (c) d18 O data from ODP Site 1090 (Billups et al., 2004), and
the data of this study from Hole 926B. (d) Nannofossil biostratigraphy from Holes 929A (Raffi, 1999; Shackleton et al., 2000) and 926B (Fornaciari,
1996). (e) d13 C data from ODP Site 1090 (Billups et al., 2004), and the data of this study from Hole 926B, compared against previously published data
from Hole 929A (Zachos et al., 2001b). (f) Magnetic reversal data from ODP Site 1090 (Channell et al., 2003; Billups et al., 2004). The time scales for ODP
Legs 177 and 154 were obtained independently. The O/M boundary position can be determined by the oxygen isotope excursion, the older end of
magnetochron C6Cn.2n (found at ODP 1090, and correlated to Leg 154 via isotope stratigraphy), and the nannofossil biostratigraphy (just above the last
appearance of S. delphix). The older end of C6Cn.2n has been assigned an age of 23.03 Ma. The shaded area corresponds to the extent of the Mi–1 event,
as plotted in Zachos et al. (2001b). The vertical grey line marks the position of the Oligocene/Miocene boundary based on the base of the magnetochron
C6Cn.2n from ODP Site 1090, which correlates well to ODP Site 926 through the isotope stratigraphy.
3.4. Evaluation of amplitude modulations
Our new age model suggests a re-evaluation of the
astronomical signature in the geological data as encoded by
the amplitude modulation terms of obliquity and eccentricity (Laskar, 1999; Shackleton et al., 1999; Pälike et al.,
2004). The results of complex demodulation at the
obliquity frequency are presented in Fig. 6(a). This analysis
shows that the 1:2 Ma amplitude modulation of the
41 ka obliquity signal is present, and approximately in
phase with the astronomical solution, for all four time
series analysed. The results here also confirm the analysis
by Pälike et al. (2004). We also computed the amplitude
variation at the climatic precession frequency, which is
modulated by the short and long eccentricity, as well as a
longer 2:4 Ma long term (Laskar, 1999). Here, we
extracted the 405 ka amplitude modulation term, as
presented in Fig. 6(b). Apart from an interval around
20 Ma the data modulation is again in phase with the
astronomical modulation, and the data sets also suggest the
presence of the longer multi-million year modulation
between 21:4 and 23.8 Ma. We did not analyse the
climatic precession frequency for the stable carbon isotope
data as this signal is not significantly above the noise
background (Fig. 3(a)), presumably due to the low-pass
filtering nature of the oceanic carbon reservoir.
3.5. Enhanced variability of carbon isotope data prior to ca.
25.75 Ma
The carbon isotope data show enhanced variability
before around 25.75 Ma (Figs. 2 and 4). At first sight, these
data look like a larger noise contribution, and their local
nature is suggested by the absence of this larger variability
in records from the Pacific (Pälike et al., unpublished data).
However, on closer inspection the variability cannot simply
be explained by individual data point outliers. Instead, the
enhanced variability appears to closely follow an astronomical forcing (Fig. 7). The analysis shows a strong
obliquity imprint in the 41% d13 C variations. In contrast
to the remainder of the data set, here the phase relationship
between carbon isotopes and astronomical obliquity is
reversed, such that light carbon isotope values correspond
to low obliquity and insolation. There are two possible
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explanations for the low values: (1) a ‘‘Mackensen’’
(Mackensen et al., 1993) effect possibly associated with
more frequent algal blooms, or (2) systematic shifts in
Cibicidoides species abundances towards those with more
depleted values. In this interval, because of the lack of C.
mundulus, species were mixed to generate sufficient
carbonate for isotope analysis. Either of these artefacts
could be related to local responses to changes in
astronomical forcing (i.e., the strength of the trade winds).
In addition, d13 C minima coincide with coarse fraction
minima over this interval, indicating the possibility of an
enhanced influx of corrosive water.
1
0
-1
3
2
-3
1
0
-1
-2
(a)
Astronomy (ETP)
SusRef
coarse fraction
δ13 C
Obliquity demodulation
Data (standardised)
Obliquity demodulation
Astronomy (standardised)
3
2
-2
9
-3
δ18 O
4. Discussion
1
0
-1
3
-2
2
-3
1
0
-1
-2
24
25
-3
26
Age (Ma)
ETP
clim. prec. filter
3
2
1
0
-1
-2
-3
0.5
13
0.0
δ C
ETP
Fig. 6. Comparison of amplitude modulation between astronomical
solution and ODP 154 data: (a) obliquity amplitude complex demodulation (Bloomfield, 1976) implemented using the software ‘‘TRIP’’
(Gastineau and Laskar, 2005). The data and astronomical solution were
demodulated at a period of 40:17 ka, and from the demodulation the
1:2 Ma amplitude modulation was extracted, and normalised to standard
deviation and zero mean. (b) Climatic precession/eccentricity amplitude
complex demodulation. The data and astronomical solution were
demodulated at a period of 18:8 ka, and from the demodulation the
400 ka amplitude modulation was extracted, and normalised to standard
deviation and zero mean.
-0.5
-1.0
0.4
0.2
0
-0.2
-0.4
13
(b)
23
δ C
22
clim. prec. filter
21
13
20
We previously suggested that the Mi–1 glaciation event
at 23 Ma was related to a rare confluence of reduced
obliquity amplitude variations (a obliquity ‘‘node’’) in
combination with low eccentricity (Zachos et al., 2001b).
The new data presented here as well as the new
astronomical solution used allow us to refine our previous
findings. First, in the new astronomical solution, there is
still an obliquity node at 23 Ma, however, it is not as
pronounced as it was in the previous astronomical
calculation of Laskar et al. (1993). In the new solution of
Laskar et al. (2004), the most extreme obliquity nodes
occur at around 19.7, 20.7, 24.4 and 25.4 Ma (Fig. 2). Our
age model is sufficiently robust to exclude the possibility
that the age for Mi–1 has to be re-assigned to either 24.4 or
20.7 Ma. At most obliquity nodes (but not all) we do
indeed observe isotope excursions toward more heavier
values in d13 C and d18 O, however, our data suggest that the
obliquity filter
19
4.1. Glaciation events and orbital confluences
Precession demodulation
Data (standardised)
Precession demodulation
Astronomy (standardised)
3
2
1
0
-1
-2
25.6
25.8
26
26.2
δ C
ETP
obliquity filter
2
26.4
Age (Ma)
Fig. 7. Close-up view of benthic d13 C between 25.5 and 26.5 Ma. Top panel shows carbon isotope data (dashed green) and astronomical ‘‘ETP’’ curve for
the interval that shows unusually strong carbon isotope fluctuations (see Figs. 1 and 2). Middle panel shows these two time series after applying a gaussian
bandpass filter centred around the climatic precession frequencies ð47 9 Ma1 Þ. Bottom panel shows these two time series after applying a gaussian
bandpass filter centred on the main obliquity frequency ð25 5 Ma1 Þ. Results suggest that strong local d13 C fluctuations are at least partly driven by a
local process.
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10
as a local rather than global feature, as judged by the much
weaker recording of this signal in the stable isotope
records. As argued by Zachos et al. (1997), the coarse
fraction signal can be influenced by three factors: (1)
changes in carbonate production, foraminiferal size variations and or flux rates of coccolithophores related to the
intensity of the trade winds; (2) changes in dilution of
coarse fraction by varying influxes of clays from the near
Amazon (Shackleton and Crowhurst, 1997) and (3) varying
amounts of dissolution. Zachos et al. (1997) argued for (3),
as they observed the general association of a small coarse
fraction with increased evidence of fragmentation, and
lower carbonate content as indicated by higher susceptibility measurements (see Fig. 3(f)). While the strong
additional signal of obliquity in the coarse fraction record
could suggest that at least part of the signal is contributed
by (globally) increased circulation of corrosive bottom
waters, the phase difference between d18 O and coarse
fraction is significant enough to suggest that this was not
the only contributing factor. Further studies will be needed
to de-convolve the relative contributions of the three
possibilities above.
link between astronomy and climatic excursions is more
intricate than previously suggested, and agrees well with
the finding by Wade and Pälike (2004) from an earlier part
of the Oligocene that glacial conditions are most readily
observed when the climate system is pre-conditioned by
low obliquity amplitude variations, and then triggered by
individual 110 and 405 ka eccentricity extremes. This
relationship is now also found for the main mid-Miocene
Mi–3b cooling event (Abels et al., 2005; Holbourn et al.,
2005). Existing and new data from the Pacific suggest that
the entire Oligocene was globally dominated by a strong
405 ka cycle, particularly in the benthic isotope record,
and that glaciations are typically enhanced during most
obliquity nodes. For example, the extreme glaciation event
Oi–1, during long 405 ka eccentricity cycle 84 OlC 13n, was
also pre-conditioned by an obliquity node (Coxall et al.,
2005). Further modeling studies are in progress that will
elucidate the link between astronomical forcing of the
oceanic carbon cycle and the carbon and oxygen benthic
isotope records. Support for the view that the long-term
obliquity amplitude variations have a profound influence
on the climate system is provided by the observation that
the strength of coarse fraction variations co-evolves with
the obliquity amplitude(Paul et al., 2000, our Fig. 6(a)).
4.3. Inter-site comparison of ODP 926 and 929 coarse
fraction data
4.2. Evolution of coarse fraction record and astronomical
forcing
A previous study (Paul et al., 2000) published a coarse
fraction record ð465 mmÞ for Site 929, which can be
compared to the record from the shallower Site 926
(Fig. 8). The deeper Site 929 shows a reduced coarse
fraction content, compatible with increased effects of
dissolution and fragmentation of foraminifera with depth.
For Site 929 in particular, there is a marked decrease in
coarse fraction for sample ages younger than 22 Ma. We
observe that individual coarse fraction maxima are
20
10
15
5
10
0
5
(%)
∆ 926-929 >65µm
0
24
20
23
10
22
Obliquity
(°)
15
Site 929 >65µm
(%)
20
(%)
Site 926 >65µm
An interesting observation is that our coarse fraction
record from Hole 926B shows a much stronger response at
the climatic precession frequencies than any of the other
data sets (Fig. 3(d)), although the signal at eccentricity
frequencies, which should modulate the strength of the
precession cycle, is much weaker. We interpret the strength
of the climatic precession cycle in the coarse fraction record
0
-10
21
22
23
24
25
Age (Ma)
Fig. 8. Comparison of previously published coarse fraction record from Site 929 (Paul et al., 2000) and the record from the shallower Site 926. Also shown
are Earth’s obliquity and the difference in 465 mm % coarse fraction between the two sites. The record from the deeper Site 929 shows a smaller coarse
fraction, as expected if this process was dissolution controlled. Both coarse fraction records are in phase at obliquity periods. Larger differences between
the two records are also in phase with high obliquity values. In addition, there appears a discernible trend of enhanced coarse fraction amplitude at
1:2 Ma obliquity amplitude maxima, as well as an increase of coarse fraction site-to-site differences at 1:2 Ma obliquity amplitude minima. Vertical
dashed lines are inserted to aid comparison of the phasing of individual peaks. All plotted data are interpolated at 5 ka time steps.
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approximately in phase with obliquity maxima, consistent
with the observation that glacial maxima are correlated
with obliquity maxima, and with a higher carbonate ion
concentration and deeper lysocline. We also calculated the
difference between the age interpolated records (Site 926
minus Site 929 coarse fraction). As a general trend, we
observe that this indicator of depth-dependent dissolution
(i.e., a deeper lysocline) decreases during maxima of the
1:2 Ma long obliquity amplitude modulation. In the
upper half of the record, we also observe that the obliquity
cycles manifested in the coarse fraction records from the
two sites have not cancelled out after taking the difference,
partly due to coarse fraction that is close to zero for
Site 929.
5. Summary
The extension of the original data set from Zachos et al.
(1997) presented here, in combination with a refined
astronomically calibrated age model, provides a standard
reference section for a time during Earth’s history where
periodic glacial advances and retreats on Antarctica were
extremely sensitive to variations in the boundary conditions, such as astronomical insolation forcing. The records
presented here provide a degree of chronostratigraphic
control around the Oligocene/Miocene transition that was
previously exclusive to the Quaternary. Pioneering the
combination of astronomical age models by N. J.
Shackleton with very high-resolution stable isotope measurements is one of the strongest contributions and turning
points in the Earth System Science that only now are we
beginning to exploit to its full extent.
Acknowledgements
This paper is dedicated to Professor Sir Nicholas J.
Shackleton, FRS, for his pioneering and seminal work in
developing an astronomically calibrated geological time
scale, and for his continuous enthusiastic support to our
work. We thank Brandon Murphy, Appy Sluijs, Jason
Newton and Robert Becker (deceased) for technical
support. This research used samples and data provided
by the Ocean Drilling Program (ODP). ODP (now IODP)
is sponsored by the US National Science Foundation
(NSF) and participating countries under the management
of Joint Oceanographic Institutions (JOI), Inc. Supported
by the National Science Foundation (EAR-9725789,
OCE-0117532). We thank B. Flower and an anonymous
reviewer for improving a previous version of this manuscript. We thank Isabella Raffi for suplying biostratigraphic data.
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