1
Geomorphological and geochemical evidence (sediment 230Th anomalies)
for cross-equatorial currents in the central Pacific
Neil C. Mitchell1 and John M. Huthnance2
1
School of Earth, Atmospheric and Planetary Sciences, University of Manchester,
Manchester M13 9PL, UK
2National
Oceanography Centre, Joseph Proudman building, 6 Brownlow Street,
Liverpool L3 5DA, UK
For correspondence: neil.mitchell@manchester.ac.uk
Keywords: sediment focusing, bottom Ekman layer, bottom boundary layer,
sediment furrows
Abstract
Shallow broad elongated sediment depressions and ridges are revealed in
multibeam echo-sounder data collected over the carbonate ooze in the central
equatorial Pacific. These features, otherwise called "furrows", have orientations
that appear locally distorted by seabed topography as expected of contourtrending currents but at regional scale typically cross contours at high angles. In
places, complex patterns suggest that formative currents have a strong timevarying component. From direction indicators, the movement of bottom waters
is north to south on average, though with some movement locally south to north.
There is a modest 18˚ average change in orientation crossing from north to south
of the equator, with features to the south oriented clockwise of those to the
north. This is as expected for a partly developed bottom Ekman layer, with
currents in the layer deflected by the Coriolis effect with opposing senses either
side of the equator. The features are less prominent on and immediately south of
the equator. We evaluated these observations along with reported 230Th
accumulation rates in sediment cores., which are curiously enhanced along the
2
equator, an observation that has been previously interpreted as suggesting
transport of 230Th bound to particles to the equator.
Limited current meter and other data and physical oceanographic models
help to explain these observations. Data from current meters 1˚ north of the
equator show a highly asymmetric mesoscale eddy motion here, aligned with the
furrows. Phase relationships between near-bed and upper ocean currents
suggest an indirect coupling of upper-ocean eddies with the lower ocean. The
bottom Ekman layer is predicted theoretically to thicken towards the equator.
The resulting reduced bed shear stress may explain the 230Th deposition and
more weakly developed furrows at the equator. Given evidence that equatorial
accumulation rates of 230Th and extraterrestrial 3He both fluctuated over the
Late Pleistocene, we explore how the ideas presented here could help to explain
how the geochemical anomalies relate to physical oceanographic processes.
Introduction
Multibeam echo-sounder data reveal abundant elongate sedimentary
features, which are evidence for a widespread benthic current in the equatorial
Pacific. The current is at a high angle to the expected geostrophic flow and, along
with Ekman layer theory, may help to explain a previously reported enhanced
deposition of 230Th and 3He along the equator.
In the central Pacific, bottom water temperatures increase eastwards
implying a general eastwards movement of ocean waters (Mantyla, 1975), which
is predicted by geostrophic calculations (Reid, 1997). On the other hand, 14C
ages of water generally increase northwards in the Pacific Ocean (Key et al.,
2004; Matsumoto, 2007) though they can be modelled with a bottom water
movement from a southern ocean source flowing as a western boundary current
northwards and then eastwards into the central Pacific (Roussenov et al., 2004).
These and geochemical tracers are scalar properties and, as the numerical
modelling of Roussenov et al. (2004) illustrates, they can be explained by
complex flow paths and their values are complicated by eddy diffusion, which is
poorly characterized in the deep ocean. Vector measurements (with current
meters and floats) recording the full time-varying movements of the ocean are
preferred but are limited in the abyssal central Pacific. For understanding
3
particle transport, as entrainment and deposition involve threshold shear
stresses, the full temporal history of water movement and in particular peak
currents are important rather than merely the mean current.
The sediments of the central and eastern equatorial Pacific have been
studied extensively over the past few decades from samples recovered by
scientific drilling and sediment coring because of potential links between their
properties and climate. For example, equatorial upwelling is linked to wind
stress driving surface Ekman currents and the upwelling in turn may affect
pelagic productivity, implying a link, albeit complex, between wind strength and
the production rate of particles (e.g., Mayer et al., 1992a). But such work is
potentially compromised if particles move significantly with abyssal currents
before depositing. Accumulation rates of particles are related to the shear stress
near the seabed amongst other factors (e.g., McCave and Swift (1976)).
Therefore, where temporal variations in depositional fluxes in
paleoceanographic studies have been used to investigate past chemical
properties of the ocean or upper-ocean properties (Beaufort et al., 2001; Chugh
and Bhattacharji; Griffith et al., 2010; Lyle, 2003; Mitchell et al., 2003; Pälike et
al., 2012), those fluxes may also have been affected by changing strength of
abyssal currents. Spatial variations in the properties of deposited particles
(Boltovskoy, 1991; Murray et al., 2000) may also not necessarily relate solely to
variations in their production or input to the ocean if there is significant
movement.
Geomorphologic features on the ocean floor can provide clues to water
and sediment movements, which may in turn help to guide the design of future
oceanographic fieldwork as well as indicate where uncertainties in
paleoceanographic studies may lie. The sedimentary features described here are
elongated ridges, which are essentially sediment drifts, and depressions, which
are similar to features termed elsewhere as furrows (Dyer, 1970). Although not
an ideal descriptor, we use the term "furrow" here for consistency. Furrows in
abyssal environments are typically 1-20 m deep sedimentary troughs extending
parallel to the flow direction for several kilometres (Hollister et al., 1974),
produced by helical secondary flow near the bed (Flood, 1983; Lonsdale and
Spiess, 1977; Viekman et al., 1992). Examples have been reported in calcareous
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pelagic sediment, such near the Samoan Passage, where they were found to be 1
m deep and spaced at 30 m (Lonsdale, 1981; Lonsdale and Spiess, 1977).
Lonsdale and co-workers (Lonsdale, 1976; Lonsdale, 1977; Lonsdale, 1981;
Lonsdale and Malfait, 1974) have characterized near-bottom flow patterns in the
southwest and east Pacific using these features. Both steady and oscillating
currents can produce furrows (Flood, 1983), which indicate the general
orientation of flow.
The earlier studies of these low-relief features relied on sidescan sonar
data collected from instruments close to the bed to generate acoustic shadows
and strong acoustic backscattering contrasts. The early 16-beam SeaBeam
multibeam sonars were noisy (de Moustier and Kleinrock, 1986) so only the
largest erosional features could be interpreted from their data (Shipley et al.,
1985). Over the last decade or two, however, the multibeam sonar technology
has improved significantly. The data acquired with them have better noise
characteristics and effective spatial resolution. The context provided by their
wider swaths helps the interpreter discriminate these current-generated
features from other elongated features, such as abyssal hills.
The furrows described here are generally more widely spaced than those
mentioned in the earlier publications. This difference may merely be an artifact
of differing sonar resolution effectively filtering a continuum of bedform scale
(the multibeam data shown here have a spatial resolution of ~112 m whereas
the Deep Tow data used by Lonsdale and others had much higher resolution).
Indeed high-resolution sonar images of such features elsewhere suggest a range
of spacings can occur in practice (Cochonat et al., 1989). Alternatively, this
difference might have a fluid dynamical origin. Current measurements in Lake
Superior showed the secondary circulation extending to a height similar to the
furrow spacing (Viekman et al., 1992). The difference may therefore relate to
different boundary layer thicknesses, though we have insufficient data to
confirm this here. These features are also morphologically similar to shallow
gullies observed in upper continental slope settings produced by density
currents (Izumi, 2004; Puig et al., 2008). Those off southern California, for
example, have spacings of 100-500 m (Field et al., 1999; Mitchell and Huthnance,
2007), overlapping with spacings of the furrows shown here.
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Sediment geochemical data have also been interpreted as evidence for
transport of particles by bottom currents to and their deposition on the Pacific
equator (Marcantonio et al., 1996), and attempts have been made to use such
data to correct depositional fluxes for particle redistribution (Francois et al.,
2004). We re-examine those arguments in the light of the new data. 3He and
230Th
accumulation rates in equatorial Pacific sediments in places deviate from
their introduction to the water column (3He from interplanetary particle input to
the ocean (Marcantonio et al., 1996) and 230Th from 234U decay (Higgins et al.,
1999; Marcantonio et al., 2001)). Thorium is strongly particle-reactive in water,
so 230Th produced by radioactive decay is readily absorbed onto falling particles.
Consequently, it has an average residence time in the oceans of only 20 years
(Henderson and Anderson, 2003). Uranium, in contrast, is well mixed and
production of 230Th from radioactive decay is predictable from water depth and
the uranium concentration. The ratio of accumulation rates of 230Th to the
expected production of 230Th in the immediately overlying water column then
suggests whether sediment has been advected into or away from a site. Further
measurements of this ratio (“Y”) termed the "focusing factor" compiled by
Kienast et al. (2007), which are shown in Figure 1a, confirm this apparent
anomalous accumulation of particles along the equator.
This simple interpretation of the 230Th data has been controversial,
however. Some authors (Broecker, 2008; Lyle et al., 2005; Thomas et al., 2000)
have argued that focusing factors are enhanced at the equator because of
enhanced scavenging by falling particles in this high pelagic productivity region.
Removal of thorium from the equatorial ocean is then compensated by spreading
of thorium by eddy diffusion along isopycnals from outside the equatorial region.
Siddall et al. (2008), who described this as a particle flux effect (PFE), extended a
global circulation model to study it and argued that it was insufficient to explain
the magnitude of observed equatorial focusing factors. Thorium is preferentially
absorbed onto finer lithogenic particles (in particular, clays), which have larger
effective surface area for a given mass (Geibert and Usbeck, 2004; Kretschmer et
al., 2010; McGee et al., 2010; Roy-Barman et al., 2005). Their smaller terminal
velocities allow them to be more easily advected laterally than larger particles,
so varied focusing factors may partly reflect varied transport of fines. For a
6
discussion of other potential issues with this method, see Lyle et al. (2007) and
Francois et al. (2007). We return to the particle flux and particle size effects in
the discussion.
In the present study, evidence for equatorial bottom currents is examined
in multibeam echo-sounder and chirp sediment profiler records collected during
cruise AMAT03RR of the RV Revelle (Dubois and Mitchell, 2012; Tominaga et al.,
2011). Further information on the flow is provided by limited current meter
datasets. The discussion examines the origins of the geomorphologic features,
the nature of time-varying currents crossing the equator here, the elevated 230Th
anomalies reported previously (Kienast et al., 2007) and theoretical reasons why
sediment accumulation may be enhanced along the equator, as well as general
consequences for paleoceanographic studies.
Datasets
Geophysical data
Bathymetry data were collected on Revelle using a Kongsberg-Simrad
EM120 multibeam echo-sounder during site survey cruise AMAT03RR for
Integrated Ocean Drilling Program (IODP) expedition 319/320 (Pälike et al.,
2010). Data were acquired in small grid surveys over each IODP site as well as in
transit, while simultaneously running seismic reflection, chirp sediment profiler
and other scientific instruments.
The EM120 multibeam sonar on Revelle has 191 beams spread over a
150˚ sector. Resolution fore-aft is dictated by the 1˚ transmit acoustic beam
widths (~70 m in 4000 m of water at nadir, widening with slant range away from
nadir). The system uses the split-beam method of bottom echo detection,
whereby the time of the echo arriving at the beam centre is determined from a
series of differences of acoustic phase detected between two sections of the
receive transducers (Hammerstad et al., 1991). Resolution athwartships
therefore depends on the time window and method used for that determination,
but is probably similar to the fore-aft resolution by inspection of the resulting
bathymetry data. Because noise is reduced by filtering the data, resolution also
in practice depends on the noise characteristics of the integrated system
including motion sensors and acoustic noise arising from sea conditions (Schmitt
7
et al., 2008). Data collected during trials with the ship stationary were found to
have a standard deviation of less than 0.2% of water depth (<8 m in 4000 m
depth) across the swath to 60˚ off-nadir (de Moustier, 2001).
The sounding data were processed using standard procedures within the
MB-System software (Caress and Chayes, 1996), including manually removing
acoustic outliers. The data were binned in 0.001˚ cells (a step that reduces noise
somewhat by averaging ~2-3 values in each cell) and gridded at 0.0005˚
resolution using a surface-fitting program (Smith and Wessel, 1990). The
effective spatial resolution of the following images is therefore ~112 m. The
shaded relief images in Figures 2-4 illustrate the quality of the data. Artifacts
include enhanced variability in the outer swath produced by noise probably from
the water column multiple of the transmitted pulse, some across-track striping
caused by noise from the vessel passing through waves and minor remaining
individual sounding outliers. An along-track striping can also be observed at the
swath nadir (such as in the centre of Figure 3b), where the system uses a
different seabed echo detection method (Mitchell, 1996). These artifacts are all
readily identifiable and can be avoided in interpretation. If not acoustically
isolated, interference from other sonars transmitting with different cycle periods
might lead to artifacts trending obliquely to the swath, but in areas of flat seabed
no such artifacts are evident in these data.
The data were analyzed with the GeoMapApp software
(www.geomapapp.org) along with other wide-swath multibeam sonar data that
have been incorporated within the Global Multi-Resolution Topography
Synthesis (Ryan et al., 2009). Different artificial sun illumination directions were
used to assess trends, using low inclination illumination to enhance subtle
features not so easily interpreted in contour maps (Edwards et al., 1984). Such
shading calculations typically involve a directional derivative of the topography,
which enhances small features relative to large features if appropriately
oriented. The shaded-relief imagery in Figures 2-4 was produced with the GMT
software system (Wessel and Smith, 1991) with the illumination directions given
in the figure captions. Those directions were chosen to show best the seabed
features we interpret here. Readers can further study these data by viewing
them when available from the National Geophysical Data Center
8
(www.ngdc.noaa.gov/mgg). Contour maps of Figures 2-4 are also provided in an
electronic supplement to this paper (link). Although contours are noisy from
across- and along-track artifacts mentioned above, deviations of contours in the
maps can be used to assess relief of the larger furrows.
{Note to editors/typesetters: text "when available" can be removed if the data
have been uploaded to NGDC by the time of publication}
Current meter data and tidal models
No current meter records are available from the centre of the region
studied here, but data from adjacent sites provide some sense of the character of
the currents. Individual near-bed current meter measurements are shown as
point clouds in Figure 5 for the sites marked MS, MC and MCD in Figure 1a
(Manganese Nodule Project (MANOP) sites S, C and CD, respectively). The data
in (a), (b), (d) and (e) of Figure 5 were provided by the Buoys Group of the
Oregon State University (http://cmrecords.net/osu/pacific/pacific.htm),
whereas the record in (c) was reproduced from Mayer (1981). (The magnetic
compass data from the latter site may have been biased (Gardner et al., 1984) so
only the dispersion of directions and speed can be interpreted from this record.)
Figure 5f shows a time-series of current speed recorded at site MANOP Site C1
(20 mab) to illustrate the influence of tides in a relatively quiescent part of the
record. Measurement durations, given in Table 1, vary over 7-14 months.
Progressive vector diagrams (PVDs) computed from the MANOP C1, C2 and C3
data are shown in Figure 6c (solid circles are shown at weekly intervals). A local
bathymetry map and the temperature records from C1 are shown in Figures 6a
and 6b, respectively. The PVDs can be contrasted with diagrams for the more
northerly MANOP-S sites shown in Figure 7a.
Further current measurements have been made north of the Clipperton
Fracture Zone (Morgan et al. (1999) reported flow data compiled by Demidova et
al. (1996)). However, the plots in Figure 6c illustrate that measurements made
over much less than a year are unlikely to report the flow properties accurately
given the variations of various time-scales including greater-than-annual periods
affecting bottom waters here, so we have shown only one of those recordings in
Figure 1a (site B, 197 days, 30 mab). The mean current of each set of current
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meter data is shown by vectors in Figure 1a (red and purple for >100 mab and
<100 mab, respectively) and is reproduced in Table 1. The vector “MM”
represents the mean current at MANOP Site M (200 mab, 13 months) from data
in Lyle et al. (1985).
Theoretical estimates of the barotropic tidal currents were made at the
locations shown by cross symbols in Figure 1a and at the MANOP C1 and CD
current meter sites. Calculations were carried out using the Oregon State
University TPX06.2 global tide model (Egbert and Erofeeva, 2002) using
MatlabTM Tide Model Driver software. Figure 8 shows scatter plots of the results,
each calculated for one lunar month from 1/1/1992 with all main tidal
constituents (M2, S2, N2, K2, K1, O1, P1, Q1, MV and MM).
Geochemical data
The 230Th accumulation anomaly or focusing factor Y is commonly defined as
(Suman and Bacon, 1989):
ò
Y=
éë Thex0 ùû rb dr
b H(t2 - t1 )
r 2 230
r1
(1),
where [230Thex,00] is the decay-corrected concentration of 230Th in sediments
corrected also for in-situ production, b is the sediment dry bulk density and  is
the production rate of 230Th in the overlying water column of depth H. The
calculation is carried out over the sediment depth range r1 to r2 corresponding to
the sediment age range t1 to t2.
The colour-filled circles in Figures 1a and 1b (with scales to right) show the Yvalues that have been reported in the literature for the Holocene and Last Glacial
Maximum (LGM), respectively. These values are from a compilation of Kienast et
al. (2007, their Table 4), supplemented with five values for the LGM calculated
from data of McGee et al. (2007). Figure 9a shows the Y-values for the region
109˚W to 145˚W plotted versus latitude (solid and open circles represent
Holocene and LGM values, respectively). Also shown in Figure 9a, large greyfilled and open circles connected by solid lines are average Y-values (Holocene
and LGM, respectively) each computed over 1˚ of latitude centred on 0˚ and 1˚N.
As explored by Kienast et al. (2007), uncertainties in individual Y-values are
commonly dominated by dating uncertainties and can be difficult to quantify.
10
Skinner and Shackleton (2005) documented a 2 ky lag of the Marine Isotope
Stage 1/2 boundary in a Pacific core compared with the glacioeustatic (global)
curve. This error would bias (increase) Y-values by 15% and may largely explain
why the LGM values, which have not been corrected for this effect (Kienast, pers.
comm., 2013), are larger than the Holocene values in Figure 9a. Kienast et al.
(2007) estimated a further 1.5 ky uncertainty could arise from difficulty in
identifying the Stage 2/3 boundary. McGee et al. (2007) also estimated 1
uncertainties in Holocene Y-values by combining uncertainties in sediment age
and density. For each of these studies (Kienast et al., 2007; McGee et al., 2007),
we computed the mean ratio of their uncertainty to value of Y in order to
construct the "typical" uncertainty bars shown in the left of Figure 9a for Y=1 and
Y=2.
A further uncertainty (of interpretation rather than values) could arise from
local transport of sediments, for example, because of a tendency for sediment to
infill basins between abyssal hills (Dubois and Mitchell, 2012; Laguros and
Shipley, 1989; Mitchell, 1995; Mitchell et al., 1998; Tominaga et al., 2011) or be
removed from basement highs or areas of scour around outcrops (Mitchell,
1995; Mitchell, 1998b). To assess this effect, the core locations (Kienast et al.,
2007; McGee et al., 2007) were read into the Geomapapp software so that the
physiography of each site could be evaluated using either multibeam bathymetry
data if available (Ryan et al., 2009) or otherwise satellite gravity-derived
bathymetry capable of resolving larger features (Smith and Sandwell, 1997). The
following core site names are from Kienast et al (2007) and McGee et al. (2007).
Two sites were found to be close to seamounts of >1000 m elevation
(TT013/MC34 and TT013/MC19) and one adjacent to a 600 m hill
(TT013/MC69). Three sites lie on or near the edge of an abyssal hill where
winnowing might be more pronounced (MANOP/B18, ODP852A and ODP853B).
The Y-values of these sites are marked in Figure 9a with cross-symbols. These
potentially biased values appear to have no particular effect on the overall trend,
however; the smaller circles connected by dashed lines are 1˚-averages after
removing these values. Because of potential enhanced scavenging associated
with hydrothermal activity (Singh et al., 2013) as well as these effects of
topography, the sites closer to the East Pacific Rise were also excluded from the
11
analysis. With these various selections, the 1˚ averages were computed from
only 3 values at 0˚N and 6 values at 1˚N for both the remaining Holocene and
LGM data. The standard errors of the means after this averaging will be smaller
than the uncertainty bars in Figure 9a and the equatorial tendency should
therefore be resolved. We return to the uncertainty issues in the Discussion.
For comparison with the 230Th results, linear sedimentation rates and mass
accumulation rates were obtained from the line of ODP drill sites at 110˚W,
where a relatively uniform carbonate content implies uniform extent of
dissolution or dilution by silica-bearing species (mean carbonate contents
computed from the full sediment core depths at Sites 849, 850 and 851 at 0.2˚,
1.3˚ and 2.8˚N (Mayer et al., 1992b) are 71.0, 68.2 and 72.0%, respectively). The
data in Figure 9c are linear sedimentation rates computed using the highresolution timescale of Shackleton et al. (1995) obtained by orbital tuning.
Values are normalized to the peak value of each period, which consistently lies at
the near-equatorial Site 849. In Figure 9d, a similar graph is shown with values
converted to mass accumulation rates using the core index properties (Mayer et
al., 1992b). The bold line in each case represents the 0-100 ka interval.
Observations
Geomorphology
The furrows and drifts can potentially be confused with topography of the
sediment surface mimicking the underlying abyssal hills. To help discriminate
them, abyssal hill trends were predicted using the orientations of sea-floor
spreading magnetic isochrons (Müller et al., 1997) and interpreted from chains
of pits caused by hydrothermal expulsions from basement (Bekins et al., 2007;
Michaud et al., 2005; Moore et al., 2007). The white dashed lines in Figures 2-4
show our interpreted abyssal hill trends. Abyssal hills produced at mid-ocean
ridges also have a rugged morphology of superimposed volcanic, tectonic and
mass-wasting features so associated topographic trends can be irregular.
Furthermore, the interpreter needs to be aware of the potential for oblique
basement trends originally produced by transform faults, propagating rifts and
other ridge discontinuities (Hey et al., 1986; Macdonald et al., 1991).
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Figure 2a shows multibeam data collected towards the eastern side of the
surveyed area. Elongated ridges and troughs seen running roughly north-south
are up to 10s of metres deep (see cross-sections in Figure 2a sampled from the
multibeam data) and cross abyssal hills (trending just west of north) at a shallow
angle. These furrows are more irregular than those described in the earlier
studies (Flood, 1983; Lonsdale and Spiess, 1977). The chirp sediment profiler
record in Figure 2b (collected between EE and E’ along the track line marked in
Figure 2a) reveals sub-bottom reflectors almost paralleling the seabed (the subbottom time to the prominent reflector at 25 ms for example was estimated to
have a coefficient of variation of 11% across the E-E' record). There is no
obvious evidence of erosion, such as truncation of reflectors at the seabed or
unconformities within the acoustic stratigraphy here. Reflectors in similar
sediment profiler records collected in the equatorial Pacific arise from acoustic
interference of reflections from impedance structures that can be complex,
varying over many scales and mostly linked to changes in density caused by
carbonate content variations (Mayer, 1980). However, the coherence of the 25
ms reflector suggests it results from a more significant impedance contrast.
Many other parts of the eastern dataset also show draping morphologies. In the
NW of the map (Figure 2a), smaller elongated features can be observed crossing
the ship track. In the chirp data (lower record in Figure 2b), they correspond
with small undulations of the seabed echo, with small diffraction hyperbolae
emanating from some troughs. However, in other parts of the chirp record
where the seabed in the multibeam data appears smooth (such as shortly after
(left of) "Course change" in E-E') we see few diffraction hyperbolae so smooth
areas probably lack furrows of significant relief (arguing against sonar resolution
being the cause of different apparent furrow spacings mentioned in the
introduction).
In Figure 3a, a series of shallower and broader troughs emanate to the
southwest from depressions of the sediment surface. (These depressions are the
pits of Moore et al. (2007) and here are aligned both parallel to abyssal hills and
perpendicular to them, possibly overlying a fracture zone.) Figure 3b shows a
fine fabric of sedimentary features trending NNE to SSW crossing the abyssal hill
13
fabric visible trending slightly west of north. Arrows mark troughs emanating
from pits towards the southwest.
Figure 3c shows two small seamounts surrounded by sedimentary moats
and an abyssal hill fabric trending ~10˚ west of north. Surrounding the two
seamounts, a fine sedimentary fabric can be observed changing trend as though
produced by flow deviated by the seamount topography. Figure 3d shows
similar features around two smaller knolls. A fabric trends slightly west of north
and appears to deviate eastwards on passing east of the knolls.
On and south of the equator, current-related sedimentary features were
generally less easily interpreted. Figure 4a shows a segment collected
immediately before the data in Figure 2 along a different heading. In this case,
the lack of apparent current-related fabric might be due to the intensity
normalization used in the shaded relief software (Wessel and Smith, 1991)
reducing contrast to compensate for the across-track striping (perpendicular to
the shading direction) but some crossing fabric can be observed ("?" in figure).
In Figure 4b, some minor finely-spaced fabrics are observed sub-parallel to
abyssal hills. Similarly in Figure 4c, fabrics run a few degrees east of abyssal hill
trends. Figure 4d shows a minor NE-SW-rending drift crossing abyssal hill
topography.
Representative orientations of sedimentary lineaments were interpreted
from the Revelle and other wide-swath sonar data, avoiding the immediate
influence of seamounts and other local physiographic features. Those
orientations were then averaged in grid cells of 1˚ X 1˚ and are displayed with
white bars in Figures 1a and 1b. The sense of flow was also interpreted from
troughs emanating from pits such as in Figures 3a and 3b and from sediment
drifts on the lee sides of features. These general directions are shown by the
black vectors in Figure 1b.
The geomorphological data suggest that bottom water flows in a general
NNE-SSW orientation. The bed features (Figures 2 and 3) are more irregular
than the furrows described previously (Flood, 1983; Lonsdale and Spiess, 1977),
suggesting that they were created by a current varying in direction. Six out of
nine sense-of-movement indicators (black vectors in Figure 1b) suggest flow to
the southwest. The presence of some indicators of northeastward flow suggests
14
that the reverse current can be locally more important, perhaps an effect of local
seabed morphology. Some orientations are west of north immediately adjacent
to the Clipperton Fracture Zone (such around 130˚W).
The chirp records acquired on Revelle were examined over the region
surveyed north of 10˚S until towards the end of the survey. The transit to the
equatorial region from the south was plagued by poor sea conditions so
reflectors are not easily interpreted there, but reflectors characteristic of the
equatorial carbonates (Mayer, 1980) become sharp within a degree or two of the
equator. Within the equatorial region, there is relatively little evidence of
erosion, such as truncation of reflectors at the modern seabed or unconformities.
This may be partly due to poor lateral resolution of the wide-beam chirp profiler,
but even fully processed (to migration stage) multichannel seismic data collected
during the Revelle cruise over IODP sites show truncation only at steep
escarpments (Expedition_320/321_Scientists, 2010b). Within the equatorial
zone, sub-bottom reflectors generally parallel the modern seabed to greater than
0.1s, as illustrated in Figure 2b.
Possible erosional truncations of reflectors become more common in the
chirp data approaching the Clipperton Fracture Zone at the northern limit of the
line at 7˚N, 125˚W. Processed multichannel seismic data from there
(Expedition_320/321_Scientists, 2010a) and short-streamer seismic collected in
transit (Dubois and Mitchell, 2012) also suggest erosional truncations on the
western Revelle transect passing "3d" in Figure 1a. Non-parallel sub-bottom
reflectors are common from this area westwards, linking with the Clipperton
Fracture Zone where erosion is well known (Johnson, 1972a; Johnson, 1972b).
The change from primarily draping recent sediments in the east to more
irregular deposition to the west is accompanied by a change in sediment type
with less carbonate content occurring with increasing depth and dissolution
(Farrell and Prell, 1991; Pälike et al., 2010; Theyer et al., 1989; van Andel et al.,
1975). Hence, the change may partly reflect changing physical properties of
sediment (grain size, density and threshold of motion, for example) making it
difficult to say whether there is a change in the properties of bottom currents
across the region.
15
Although the trend is not strong, the morphologic data show a change in
furrow orientations across the equator, e.g., changing clockwise from Figure 2 to
Figure 4d. This change can also be seen in orientations in Figure 1a. We
computed an average orientation from the digitized furrow directions for the
region east of 130˚W (were data lie on both sides of the equator) separately for
1˚ to 5˚N and 1˚ to 5˚S. The northern set has a mean of 16.0˚±1.3˚ (1) for 107
samples. Unfortunately, only six orientations were recorded in the latitude band
south of the equator but nevertheless they were consistent with surrounding
orientations observed during digitization (visually representative) so their mean
of 34˚ and minimum of 29˚ are considered reliable. These values suggest an 18˚
clockwise change in the furrow orientations occurs crossing the equator north to
south with a probable minimum change around 12˚.
Current meter data and tidal current models
The progressive vector diagrams (PVDs) in Figure 6c illustrate the
unsteadiness of the flow at sites adjacent to the equator. The deployments at 20,
42 and 170 mab (metres above bed) show translational movements lasting many
weeks to months within the lower water column as well as more nearly circular
mesoscale eddies (e.g., the circulation towards the end of the 170 mab C1 data).
The C3 42 mab PVD has an annual mean displacement of 329 km along N082.5˚E,
almost parallel to the expected geostrophic flow illustrated by the mid-water
record for C3 (top-left in Figure 6c). The upper-right PVD is the 42 mab C3
record with the annualized mean drift removed, showing that the remaining
movements are one-dimensional, rather than circular as expected for
symmetrical eddies. They have an average period of 7 weeks that is comparable
to that of mesoscale eddies. The record is not quite periodic on an annual basis.
Comparing the upper and lower water column records at C3, the net movement
after 10 weeks (first open symbol along the records) in the upper record is
opposite that in the deeper record. The following 10 weeks of movements are
more nearly in phase.
Site C1 was located adjacent to an abyssal hill (Figure 6a), which appears
to have produced different movements in the two records shown in Figure 6c.
Between 30 and 40 weeks, the PVD at 20 mab suggests a circulation, while that
16
at 170 mab suggests an overall eastward translation (flow less impeded by the
hill?). The mean velocities of the MANOP-C1 sites (Figure 1a and white bars in
Figures 5b and 5e) are roughly in line with the peak velocities in Figures 5b and
5e towards the south and southwest. The upper water column record at C2,
which started on a similar date to those at C1, shows mainly an eastward
translation expected of the geostrophic flow. Whereas the cross-equatorial
movement between the C2 and the 170 mab C1 records is in opposite phase for
the first two weeks, it is more in phase around weeks 5-10. Although the
heading at MANOP-CD is unreliable, the character of this record at 4.7 mab
seems similar to that of the C1 20 mab site, though with smaller velocities
corresponding with the lower altitude of the meter. The two temperature
records in Figure 6b are parallel, apparently recording the same bottom
temperature gradient, with no obvious change such as with varying boundary
layer structure. If a bottom mixed layer is present, it is presumably less than 20
m thick here.
The MANOP S sites were unfortunately located next to seamounts (Figure
7) but they provide records farther from the equator to compare with the C sites.
Whereas the southerly site data ("short" in Figure 7) show a general southward
or south-southwestward movement, with the deeper meter slower than the
shallower meter, a general westward or southwestward movement was
recorded at the S1 meters, with greater mean velocity at the deeper site. The
latter records also show evidence for mesoscale eddies, whereas the former do
not. The MANOP-M record (Lyle et al., 1985) also shows a translation (southsouthwestward) somewhat similar in character to the MANOP-S "short" records
though with more of a superimposed circular movement of multi-monthly
period.
The computed tidal currents at the MANOP-C site (Figure 8) are less than
2 cm s-1 and largely explain the oscillation shown in Figure 5f (vertical lines are
spaced every 12.42 hours). Farther east, tidal currents can reach 3 cm s-1 but are
nevertheless still small and lie oblique to the furrow orientations (Figure 1).
Besides those at MANOP-S, the current meter data suggest long-period
oscillatory movements broadly in line with the furrow orientations. As the
morphology is more likely to reflect the peak current than the mean current, the
17
greater peak current speed at MANOP-C within the carbonate ooze zone where
we observe these features is interesting. A more modest 12 cm s-1 peak current
speed was recorded at MANOP-M (Lyle et al., 1985), similar to that at MANOP-S
(Figure 5).
230Th
variations
Regional and local variability
The high Y-values about the equator in Figure 1a have been noted
previously (Kienast et al., 2007; Marcantonio et al., 2001; Siddall et al., 2008).
Figure 9a reveals that the anomaly extends from 0˚N to perhaps 2˚S. The LGM
values are also elevated about the equator. Adjacent Y-values can be markedly
different. Examples include around 140˚, 110˚ and 102˚W on the equator (Figure
1a). Inter-site variability is less obvious away from the equator in the lines of
data at 140˚ and 110˚W, where Y-values more nearly equal unity. Much of this
variability could be due to effects of local bathymetry on the currents and
gravity-driven sediment transport as outlined earlier.
After screening out sites potentially affected by local sediment transport
variability, the remaining Holocene values in Figure 9 still show an equatorial
signature and the mean values at 0˚ and 1˚N are little changed. The equatorial Yvalue of ~2 is derived from three remaining values and the decrease towards
Y=1 at 1˚N is derived from six remaining values. However, the elevated value at
2˚S is derived from only a single sediment core.
Although representing a longer interval of sedimentation, the upper 100
k.y. of sedimentation and mass accumulation rates along the ODP Leg 138 sites
(Figures 9b and 9c) show a similar pattern about the equator. The region 2˚S to
0˚N over which 230Th and bulk sediment accumulation is enhanced coincides
with the region where we have found the furrows and sediment drifts to be more
difficult to interpret (Figure 4).
Interpretation
The furrow pattern suggests that they were generated by a current
running mainly between NE-SW and N-S (Figure 1a). From irregularity of some
furrows (Figure 3) and the current meter data from MANOP-C (Figure 6), the
18
formative current is irregular, oscillating with periods varying from tidal to
greater than annual. The lower MANOP-C3 meter shows an annual mean drift
expected of geostrophy (Reid, 1997), though over shorter periods suggests a
movement of up to several hundred kilometres oriented parallel to the furrows.
The data from several of the other current meters, however, suggest a mean drift
towards the southwest to south, compatible with the current direction indicators
in the multibeam data, which show a preference towards the south over north
with a 2:1 ratio. At the equator, a reduced shear stress in the equator-crossing
currents (of varied periods) is interpreted from furrow morphologies in the
multibeam data (Figure 4) being less distinct than north of the equator and the
anomalous accumulation of 230Th and 3He (Marcantonio et al., 2001;
Marcantonio et al., 1996). The furrow orientations in Figure 1a do not obviously
relate simply to the bathymetry contours at large scale as would occur with a
geostrophic current. For example, northeast of the "4b" annotation the furrows
roughly parallel the 4 km contour but next to the"4d" annotation, they are
perpendicular it. A similar lack of correspondence to contours can be observed
elsewhere in Figure 1a. If a bottom Ekman spiral accompanies the NE-SW
oscillating current (Figure 6c), periodic reversal of that spiral and associated
bottom current deflection may also help to explain the furrow irregularity
compared with those in the literature.
Discussion
Origins of the furrows and their generating current
The orientations and morphologies of sedimentary bedforms tend to
reflect the more extreme currents capable of re-suspending particles or
influencing patterns of non-deposition and deposition, hence it is not surprising
that the furrow orientations bear little relation to the expected eastward
geostrophic current (Reid, 1997). That the furrows appear over the carbonate
ooze despite the low current speeds and occur less commonly in the carbonate
poor area north of the Clipperton Fracture Zone may originate from a smaller
threshold of erosion; Lonsdale and Southard (1974) found half or lower
threshold velocities of carbonate ooze compared with red clay.
19
The furrows lie oblique to the peak tidal currents in Figure 8, which are
strongest in the NW and SE quadrants, where they approach 3 cm s-1. These tidal
currents are modest relative to the following re-suspension thresholds in erosion
experiments on carbonate ooze. Southard et al. (1971) found thresholds of 7-10
cm s-1 immediately after rapid deposition and 15-20 cm s-1 after a few hours.
Nevertheless, despite their small speeds, interaction with the bed can give rise to
locally enhanced currents capable of re-suspending particles (Peine et al., 2009;
Turnewitsch et al., 2008). Furthermore, phytodetritus, which is common in
photographs of the equatorial Pacific bed (Smith et al., 1996), can be resuspended by smaller currents than inorganic particles and this may help to
explain the sediment focusing (Marcantonio et al., 2001). Expressed in terms of
shear velocity u*, thresholds of 0.4-0.8 cm s-1 were reported in laboratory flume
measurements by Beaulieu (2003) and 0.4-1.2 cm s-1 for experiments on surface
aggregates recovered from core tops by Thomsen and Gust (2000), compared
with much larger 1.37 cm s-1 for the erosion experiment on a pelagic carbonate
silt of foraminifera and coccolith particles (Southard et al., 1971). If bed shear
stress under tidal currents 0 is equal to Cd|u|2 (Taylor, 1919) with typical Cd =
0.0025 (dimensionless), the tidal shear velocities u* (=(0/)1/2) reach 0.15 cm s1.
Although small, tidal current shear stress, when added vectorally to shear
stress of other motions, contribute about 1/3 of the minimum suspension
threshold stress.
In numerical models of internal waves (Simmons, 2008), wave beams
generated by the barotropic tide crossing bathymetric features in French
Polynesia lie almost exactly parallel to these furrows, particularly towards the
west. However, the semi-diurnal movements in the MANOP-C1 near-bottom
record (Figure 5f) are readily explainable by only the barotropic tidal
movements (Figure 8). Internal tidal amplitudes deduced from satellite
altimeter data (Shriver et al., 2012) also are reduced in this equatorial region.
Nevertheless, we cannot rule out the possibility that enhanced flows may be
produced locally by interactions with bottom morphology (Turnewitsch et al.,
2008) or that the furrows developed during geological periods when internal
tide movements may have been different, for example, if lowered sea level led to
20
stronger internal wave generation. Modelling is required to investigate this
possibility.
Internal gravity waves (lee waves) are produced by steady flow across
abyssal hills and other topography. They are not resolved in global ocean models
but Scott et al. (2011) modelled their generation rates using estimates of abyssal
hill topographic properties from satellite altimetry (Goff, 2010) and output of the
HYbrid Coordinate Ocean Model (HYCOM) to establish the pattern of bottom
water long period (>tidal) flow. Figure 11 in Scott et al. (2011) shows a broad
region of lee wave generation extending from the equator northwards.
Turbulence associated with lee waves may help to explain the generally elevated
shear stresses needed to produce the furrows.
Isopycnal diffusivities estimated by Davis (2005) from float trajectories at
900 m depth at the equator exceed 8000 m2 s-1. While diffusivity major axis in
his maps is generally oriented along the equator, east of 140˚W the meridional
component becomes more significant and the major axis is, in places, aligned
with the intermediate period motion shown for the deep MANOP-C sites (Figure
6c). In the above comparisons of the middle and deep ocean meter data in Figure
6c, they are only intermittently in phase, so the lower ocean eddies are not
directly coupled with the upper ocean, indeed no such connection appears in
lowered ADCP data (Kunze et al., 2006). Nevertheless, the 7-week period of the
motion described earlier is typical of mesoscale eddies. The PVD shown in Lyle
et al. (1985) also shows a roughly bi-monthly acceleration and deceleration of
southward flow at MANOP-M. A bottom mesoscale eddy motion may therefore
be the ultimate cause of the furrows, for example, from hydrostatic pressure
variations induced by the dynamic topography of migrating upper ocean eddies,
given the alternating in- and out-of-phase motion between bottom and upper
water column current meters outlined above.
Annual mean currents
The pattern of mean bottom currents obtained by averaging the current
meter records (Figure 1a), contain some curious differences between meters
placed at different altitudes. The MANOP C3 record shows an annual eastward
movement consistent with the eastward movement expected from interpretation
21
of temperature records and geostrophy (Mantyla, 1975; Reid, 1997). More
recent numerical modeling incorporating vorticity suggests that the mean flow in
the lower water column should turn to parallel the equator here (Roussenov et
al., 2004; Williams et al., 2002).
In contrast, current meters at MANOP C1 (20 and 170 mab) and M (200
mab) show southward or southwestward flow. Unfortunately, the deployments
of C1 and C3 were over different dates, but interestingly, the lower-ocean 14C
data in Roussenov et al. (2004) show a tongue of 14C-depleted water extending
from the north into the equatorial region suggesting some southward movement.
The values at 4000 m depth in Figure 1b (from Key et al. (2004)) show this 14C
feature centred a little west of 140˚W. We speculate that a southward mean
current may occur near the bed and diapycnal diffusion of the associated 14Cdepleted water has created this anomaly.
230Th
and sediment movements implied by the Y-values
A simple observation suggests that the elevated Y-values along the
equator are unlikely to be caused by the particle flux effect (PFE) alone; the
latitude distribution of particles falling to the seabed is much broader than the
latitude range of elevated Y-values. The pattern of standing stock counts of live
foraminifera in net trawls at 140˚W (Watkins et al., 1996; Watkins et al., 1998)
are shown in Figure 10b (abundances were averaged over the depth interval 0100 m). Grey and solid symbols show trawls during unusual El Niño conditions
and more typical conditions, respectively. The solid symbols and the carbonate
fluxes in Figure 10c derived from sediment traps (Dymond and Lyle, 1994; Honjo
et al., 1995) (also see Berelson (1997)) suggest that particles fall to the seabed
over a region ~10˚ across, in contrast to the only 2˚ of enhanced 230Th
accumulation. Figure 10a shows (open circles) averages of the selected data in
Figure 9a, in particular the value at 1˚N is an average of 6 values so the
northward decrease towards Y=1 at 1˚ is well constrained.
Magnitude of the PFE
Some uncertainty over the magnitude of the PFE may nevertheless
remain. Siddall et al. (2008) attempted to address the PFE magnitude by
22
adapting an ocean circulation model (Müller et al., 2006) to account for particle
fluxes and eddy diffusion. However, water 14C contours predicted by the model
of Müller et al. (2006) differ from measurements in the central Pacific, suggesting
that advection there is not perfectly modelled and the isopycnal diffusion
parameter 1000 m2 s-1 used in the modelling was not confirmed with local
independent measurements. Indeed, high kinetic energy in the lower ocean
within output of the HYCOM (Scott et al., 2011, Figure 9) implies elevated eddy
diffusion. The meridional component of isopycnal diffusion from the upperocean floats (Davis, 2005) approaches 6000 m2 s-1 east of 140˚W and Figure 6c
suggests comparable eddy movements in the lower and upper ocean.
Lateral fluxes of 230Th
A rough assessment of the magnitudes of the 230Th and particle fluxes can
be obtained from continuity. To simplify this analysis, we assume the system is
one-dimensional (particles and 230Th are transported with latitude while zonal
fluxes at the west and east boundaries of the area are roughly equal or small).
This could be justified by the relatively coherent pattern of the seabed features
(Figure 1a) and other aspects such as primary input from the ocean surface
being aligned with the equator, but may be less easily justified in other respects
as we outline below (hence, this calculation is intended to help explore the issues
only). From continuity, the southward total flux Q of 230Th would be related to
the Y-values by:
(Y -1)H b =
¶Q
¶y
(2)
where y is latitude distance (m). Integrating with distance yields:
y
Q - Q0 = b ò (Y -1)H dy
(3)
y0
where Q0 represents the southward flux at latitude distance y0 (recognizing that
strictly speaking the system is not closed (Francois et al., 2007)).
Figure 10e shows numerical evaluations of equation (3) using the
different curves in Figure 10a (y0 corresponding with the southernmost Y-value
shown) and with =0.0265 dpm m-3 yr-1 and the averaged bathymetry in Figure
10d for H. The continuous line shows the result using the interpolated Y-values
23
in Figure 10a (i.e., the solid line). This represents the net flux from both water
column transport and from 230Th adhered to transported particles. The curve
continues increasing north of 2˚N, a result of Y-values slightly greater than unity
there.
Water-borne fluxes of 230Th
Water 230Th concentrations can be used to evaluate the possibility of
230Th
meridional fluxes arising from water advection. Figure 11 shows total
230Th
activities provided by R. Anderson of Lamont-Doherty Earth Observatory
at http://usjgofs.whoi.edu. These measurements were made along 140˚W. The
profiles in Figure 11a generally show quasi-linear trends with depth typical of
interactions with falling particles (Marchal et al., 2000). In Figure 11b values
have been interpolated to the four depths shown. Although the data represent
both particulate and dissolved 230Th, their ratio is usually low, e.g., only 0.18 in
one study at 23˚N (Roy-Barman et al., 1996), so the data represent mostly
dissolved 230Th. The activities are inversely related to vertical particle fluxes
recorded in sediment traps shown in the upper panel of Figure 11b (data from
Honjo et al. (2008)), suggesting that meridional advection is minor in the
majority of the water column, as also suggested by the mid-water current meter
data (Figure 6c). (Exceptions are the 230Th activities at 12˚S, which are curiously
low despite low particle flux.)
The deepest 230Th activities at the equator, 2˚N and 5˚N (highlighted by
grey ellipse in Figure 11a) are all anomalously high, however. Although an
explanation in terms of 230Th release by dissolution of particles at the bed
(Nozaki et al., 1981) cannot be ruled out, no such elevated values exist at the
other locations in Figure 11a. Elsewhere in oceanic gyres elevated basal values
are rare (Nozaki et al., 1981; Nozaki and Nakanishi, 1985; Roy-Barman et al.,
1996; Francois et al., 2007). We suggest that a southward benthic current is
affects the benthic anomalies here, given the southward tongue of depleted
lower-ocean 14C (Figure 1b), the southward annual mean current at two of the
three benthic meters at MANOP-C and the geomorphologic data indicating
mainly southward flow. In Figure 11b, the 230Th activity at 4000 m seems
depressed at 5˚S (i.e., does not rise south from the equator like those at 3000 and
24
3500 m), a possible indication that the benthic current persists across the
equator carrying 230Th-depleted benthic water from the equator.
The activity anomalies of these deep samples combined with the current
data can be used to estimate an effective timescale over which the 230Th activity
has a tendency to return to its normal value (this will be an upper bound given
that some effect of benthic dissolution of particles (Nozaki et al., 1981) may
occur). The meridional gradient of 230Th activity is 0.051 dmp m-3 degree-1 for
the 4000 m interval in Figure 11b from the equator to 9˚N. In Figure 12,
deviations are shown of the deepest sample values from regression lines fitted to
the remaining values. Their average deviation (0.37 dpm m-3) divided by the
meridional gradient implies around 7.1˚ of effective displacement. If the current
has an annual displacement similar to that shown by the 170 mab C1 current
meter (the PVD of this meter in Figure 6c has 2.38˚ of annual latitude
displacement), the movement would occur in 3 years. The thickness of the
current must be much smaller than the 750 m elevation above the bottom of the
3550 m deep samples given that diapycnal diffusion would otherwise affect
those sample activities also.
Fluxes of 230Th associated with particles considering the PFE
The dashed line in Figure 10e shows an estimate of the 230Th flux
associated with near-bed particle movements obtained by subtracting the Yvalues predicted by Siddall et al. (2008) to arise from the PFE from the core Yvalues (replacing Y-1 in equation (3)). For this calculation, the PFE curve shown
by the grey line in Figure 10a was reconstructed from the peak equatorial values
in their Figure 6 and the mean latitude positions of the Y=1 contours in their
Figure 5. The resulting 230Th flux shown by the dashed line in Figure 10e rises to
a maximum around 1˚N and is relatively uniform north of there. If the PFE were
larger in magnitude for the reasons given earlier, the 230Th fluxes associated with
sediment movements would be smaller, returning closer to zero north of the
equator. Note that negative Y-values are not needed to signify reduced
deposition, rather it is the Y-value relative to the value expected from the PFE,
hence the PFE is important for resolving the amount of reduced deposition off
the equator. (Without a significant PFE, it would be curious that Y-values return
25
to unity north of 1˚N if this region is the source of particles to the equatorial
sites.)
After the selection described earlier for local topographic effects, the
number of core measurements contributing to the average Y-values in Figure 10a
is small. Whereas 6 values contribute to the average at 1˚N and 3 to the
equatorial average, most remaining points represent only single values. To
illustrate the effect of uncertainty, the dotted line in Figure 10a was obtained by
subtracting 0.5 from the value at 2˚S and 1 from the two values at and north of
2˚N using the uncertainties from McGee et al. (2007) scaled to the appropriate Yvalues. Although these alterations are arbitrary (we have no information to
suggest these particular values were in error), this effect seems possible for the
small number of values. The resulting pattern of flux peaks at 1˚N and declines
farther north (dotted line in Figure 10e).
An order-of-magnitude estimate of the meridional particle flux that these
230Th
fluxes may represent can be obtained by dividing by measurements of
sediment 230Th concentrations. Typical unsupported values for Holocene
sediments are around 10 dpm g-1 in the cores studied by McGee et al. (2007) and
in cores B18 and B25 of Marcantonio et al. (2001), whereas values nearer to 5
dpm g-1 were found in PC18 and PC72 by Marcantonio et al. (1996). Using the 10
dpm g-1 value, the 230Th fluxes in Figure 10e imply particle fluxes of around 0.1 g
m-1 s-1. This is an underestimate as it ignores degradation of organic carbon and
partial dissolution of carbonates and opal that occurs before burial (Berelson et
al., 1997). Nevertheless, the value is small, as might be expected here given the
weak currents. For comparison, measurements of currents and suspended
sediment concentrations during benthic storms in the northwest Atlantic were
used to estimate fluxes reaching 0.4-1.0 g m-1 s-1 (Gross and Williams, 1991).
Fluxes of 230Th associated with fine particles
Alternatively, the preferential absorption of 230Th onto finer lithogenic
particles (Geibert and Usbeck, 2004; Kretschmer et al., 2010; McGee et al., 2010;
Roy-Barman et al., 2005) implies that the bulk sediment transport could be
smaller. The particle-solution distribution coefficient for 230Th has been
estimated to be 2-3 orders of magnitude greater for lithogenic particles than for
26
carbonate and opal (Luo and Ku, 2004). Elevated turbulence associated with lee
wave generation continues north of the Clipperton Fracture Zone (Scott et al.,
2011) where dissolution has left surface sediments with <20% carbonate
(Berger et al., 1976). Although, once deposited and compacted, erosion
thresholds of clays can be greater than those of carbonate ooze (Lonsdale and
Southard, 1974; McCave, 1984), clay particles released from aggregates by the
dissolution occurring within the water column (Jansen et al., 2002) may carry
absorbed 230Th with the mean currents southwards if kept suspended. Indeed,
bottom photos taken at MANOP-CD (Gardner et al., 1984; Thorndike et al., 1982)
and others viewable in the Geomapapp database (www.geomapapp.org) show
little turbidity. Nephelometer readings at 134˚W reached only 13 g l-1 near the
bed at the equator (Eittreim, 1984). Therefore, 230Th transport by a more dilute
suspension of fine particles is suspected. Particle concentrations at MANOP-S1
estimated from nephelometer data at 0.6 mab were 4-8 g l-1 (Gardner et al.,
1984). Such concentrations, if distributed over (for the sake of argument) 100 m
and if advected southwards at 1 cm s-1 (a typical mean current in Table 1), would
generate a mass flux of 4-8 mg m-1 s-1. This is only a small fraction, ~1%, of the
total sediment flux based on the equatorial 230Th concentrations, nevertheless, it
is comparable with the amount of deposited non-carbonate fractions. We thus
conclude from all the above that equatorial focusing of sediment is likely to be
occurring, though contributions from fine particles to meridional 230Th flux may
also be significant and need to be better characterized.
The origins of abruptness of 230Th deposition at the equator
The equatorially elevated 230Th accumulation (Y-values) occurs over a
remarkably narrow range of latitude, only ~2˚ (Figure 9a). The change implies
sharply declining bed shear stress within the lower turbulent bottom boundary
layer (BBL) in currents crossing the equator (Marcantonio et al., 2001). The
magnitude of that decline need not be as abrupt as that of the Y-values, given that
accumulation and re-suspension of sediments occur over thresholds of shear
stress. Nevertheless, given that re-suspension thresholds and flows are likely to
vary, the abruptness of the change in Figure 9a still seems remarkable. In the
27
absence of accurate in situ data, the following speculatively explores some
theoretical ideas that may help to explain this phenomenon.
The accumulation rates of particles in the oceans are often modelled as a
function of bed shear stress. For example, McCave and Swift (1976) introduced
the following expression for the mass accumulation rate R of fine particles:
æ t ö
R = Cws ç1- 0 ÷
è tc ø
(4)
where C is the particle mass concentration just outside the viscous sub-layer
near the bed, ws is their settling velocity, 0 is the bed shear stress and c is a
critical shear stress for deposition. Equation (4) suggests that a current with
particles of a given C and ws could abruptly deposit if 0 decreases below the
threshold c as we suggest below could occur because of the effect of Coriolis on
the bottom Ekman layer. C and ws potentially vary spatially, for example,
reflecting the varied particle flux from surface waters and their varied
compositions, but this is not such a concern here as we are evaluating change in
deposition relative to that expected from the primary input, i.e., how Y-values
vary. (Incidentally, an abrupt decrease in particle concentration associated with
enhanced deposition along the equator might explain why the 230Th anomalies
do not continue farther south of 2˚S (Figure 9).) The current oscillates with
periods of two months or more, such as at MANOP-C1 and -C3 (Figure 6c),
carrying particles several hundred kilometres if continually suspended, and
depositing where intersecting the equator (which thus acts like an onedirectional filter in this interpretation).
Reduced bed shear stress at the equator is predicted theoretically from
the expected dynamics of the bottom Ekman layer. The following is adapted
from arguments in Apel (1987). An interior flow up is driven by ocean pressure
gradients and balanced geostrophically (non-accelerating). The interior flow is
assumed to vary more slowly than the Coriolis frequency f. The Ekman current
components (ub, vb), where ub is aligned with the interior flow, represent the
balance of Coriolis force with the effect of eddy viscosity (friction) within the
Ekman layer. In Apel (1987), this is written:
Au ¶2 uB
(5a)
- fvB =
r ¶z 2
28
Au ¶2 vB
(5b)
fuB =
r ¶z 2
where f is the Coriolis parameter equal to 2sin ( is the Earth's angular
velocity (radians s-1) and  is latitude), Av is eddy viscosity,  is water density and
z is altitude above bottom. Equations (5a) and (5b) can be manipulated to
produce fourth-order differential equations involving ub and vb separately. With
the boundary conditions corresponding to the Ekman current decreasing to zero
at the bed and at a long distance from the bed in the interior flow, the following
are possible solutions:
uB = -uge-z cosz
(6a)
vB = uge-z sin z
(6b)
where  is z/B and B is the Ekman layer thickness equal to (2Av/f)1/2. Both
components of shear stress within the flow approaching the bed (=Au/z,
Av/z) can then be shown to equal
tu = tv =
Av ug
dB
(7).
If eddy viscosity Av were uniform, and substituting the parameters for B,
equation (7) implies that these two shear stress components are both
proportional to (sin)1/2 and decrease sharply to zero at the equator. Reducing
stress represents thickening of the Ekman layer so that the Ekman spiral occurs
over a greater depth interval and the vertical gradients in ub and vb both
decrease.
In practice, the spiral is unlikely to be fully developed. For the eddy
period of around 7 weeks observed, the assumption that flow varies more slowly
than f is strictly speaking violated within half a degree of the equator. The
assumption of uniform Av is also problematic. Nevertheless, Ekman currents
would seem to provide a simple explanation for the elevated Y-values at the
equator, less well-developed furrows in the geomorphologic data and change in
furrow orientation across the equator. Although we lack the current meter data
to confirm that Ekman spirals exist, the zigzag motion in Figure 6c for C3 (42
mab) and C1 (170 mab) is as expected of an Ekman layer deflection.
29
Implications for Paleoceanography
High glacial Y-values at equatorial sites have been noted to coincide with
evidence of enhanced flow in the deep western boundary current into the Pacific
basin from intervals of coarser sortable silt recovered near New Zealand (Hall et
al., 2001), but a physical basis for such a connection has not been provided. The
above discussion provides several mechanisms by which the general circulation
and equatorial Y-values may be linked. Turbulence associated with the lee wave
generation described by Scott et al. (2011) was likely enhanced when the flow
was stronger. If the bottom Ekman spiral is more strongly developed when the
mean flow is stronger, this should lead to a larger deviation of the bottom
current relative to the overlying flow and hence greater convergence at the
equator, suggesting that the equator may become a better sediment trap also.
The bottom Ekman explanation involves elevated bed shear stress in
currents flowing towards the equator causing non-deposition or re-suspension
of particles in areas away from the equator, combined with an abrupt decline in
shear stress at the equator. Crucially, the mechanism needs to involve low
equatorial stress in order to lead to anomalous equatorial deposition, otherwise
particles would have been merely swept across the equator during periods of
generally stronger currents. At times when the equator-crossing mean current
was strengthened relative to shorter period (oscillating) currents such as tides,
the Ekman layer should have become better developed and hence the equatorial
effect may have become greater with the greater mean current. This may help to
resolve an apparent contradiction of high accumulation rates occurring at the
equator at times when the currents were stronger in interpretations of the
geochemical data (Marcantonio et al., 2001).
Measurements of 230Th and 3He accumulation rates in equatorial cores at
140˚W have revealed fluctuations in apparent sediment focusing throughout the
450 ky of the late Pleistocene, with focusing factor abruptly rising to Y=4
principally at the onsets of glaciations and 1<Y<2 at other times (Marcantonio et
al., 2001; Marcantonio et al., 1996). At 110˚W, an equatorial core shows
persistently elevated Y-values of 2-5, with the LGM values towards the top of this
range (Loubere and Richaud, 2007). These Y-values seem too large to be
dismissed in terms of errors in sediment dates or other uncertainties (Kienast et
30
al., 2007). Uncertainties remain, however, as to the extent to which such values
represent enhanced movement of the bulk sediment to the equator or, if 230Th is
carried by finer lithogenic particles, chemical conditions leading to release of
such particles from aggregates, as well as their translation towards the equator.
They may also have been affected by enhanced isopycnal diffusion, for example,
if wind stress affecting the ocean surface was elevated, as winds have been
suggested to have been generally strengthened during glacial periods (Mayewski
et al., 1994).
Finally, we urge caution when interpreting abrupt changes in
accumulation rates from equatorial sediment cores. For example, Pälike et al.
(2012) recently claimed rapid variations in the carbonate compensation depth
(CCD) occur in the equatorial Pacific by extrapolating carbonate accumulation
rates within recently collected cores. Rapid changes will also occur at drill sites
as they drift with plate tectonics across the equator and also because of unknown
but potentially large variations with varying bottom currents.
Conclusions
The seabed morphological data presented here reveal abundant irregular
furrows, most common north of the equator. Two-thirds of current direction
indicators suggest the furrows have been created by southwestward flow.
Furrow orientations change clockwise going north to south of the equator.
These observations suggest currents flow towards the southwest but probably
with a strong oscillatory component and that they are oriented clockwise in the
south because of the change in sign of the Coriolis force on crossing the equator
affecting bottom Ekman currents.
The near-bed current meter records are generally compatible with
southward mean and peak flows. However, the records from one meter at 42
mab (MANOP-C3) show eastward annual-mean flow (compatible with
geostrophic estimates) while flows of ~7 week periods are NE-SW and parallel to
the furrows. Intriguingly, this intermediate motion is in line with the major axis
of eddy diffusivity resolved east of 140˚W in the upper ocean (900 m depth) from
float trajectories (Davis, 2005). Based on varied phase relationships between
near-bed and upper ocean current meters, some of the benthic motion creating
31
the furrows is speculatively suggested to originate from indirect coupling of
upper-ocean eddies with the lower ocean, for example through hydrostatic
pressure variations induced by the former.
In the light of the geomorphologic observations, we also re-examined the
claim of sediment focusing (Marcantonio et al., 2001). Theoretical relationships
suggest that a reduction in near-bed shear stress can occur on crossing the
equator, because the bottom Ekman layer expands at the equator. The Ekman
layer may have been better developed at times when the current of greater than
tidal periods was stronger. Therefore, the apparent paradox of high
accumulation rates on the equator (large focusing factors) during times of
stronger currents that arises from interpreting the large Y-values as representing
periods of sediment transport (Marcantonio et al., 2001) may be explained by a
strengthened Ekman layer accompanying the stronger currents. Such processes
have the potential to explain how the equator has remained an important
sediment trap for particles transported across it. In order to improve the use of
this geochemical tracer, our continuity-based calculation of 230Th flux and other
observations suggest that eddy diffusion and advection need to be better
characterized as well as contributions from fine particle transport.
Acknowledgements
For their help in collecting the AMAT03RR cruise data, we thank the captain and
crew of Roger Revelle, Mitch Lyle and the shipboard scientists. Cruise costs were
supported by grants from the NSF (OCE-9634141 to Lyle) and NERC
(NE/C508985/2). The Oregon State University tidal modelling group (Gary Egbert
and Lana Erofeeva) are thanked for online access to their TPX06.2 tidal model.
Figures in this article have been generated with the aid of the GMT software system
(Wessel and Smith, 1991). The following are thanked for email discussions and
help in locating datasets: Ric Williams, Nathalie Dubois, Stephanie Kienast, Mark
Siddall, Bob Anderson, Gideon Henderson, Larry Mayer and Roger François. Rob
Scott is thanked for providing a copy of his model output and answering
questions concerning it. We thank Bob Anderson and anonymous reviewers for
helpful suggestions, which led to a significant improvement to this paper.
32
References
Apel, J.R., 1987. Principles of ocean physics. Academic Press, New York, 634 pp.
pp.
Beaufort, L., de Garidel-Thoron, T., Mix, A.C. and Pisias, N.G., 2001. ENSO-like
forcing on oceanic primary production during the Late Pleistocene.
Science, 293: 2440-2444.
Beaulieu, S.E., 2003. Resuspension of phytodetritus from the sea floor: A
laboratory flume study. Limnol. Oceanogr., 48: 1235-1244.
Bekins, B.A., Spivack, A.J., Davis, E.E. and Mayer, L.A., 2007. Dissolution of
biogenic ooze over basement edifices in the equatorial Pacific with
implications for hydrothermal ventilation of the oceanic crust. Geology,
35: 679-682.
Berelson, W.M., Anderson, R.F., Dymond, J., DeMaster, D., Hammond, D.E., Collier,
R., Honjo, S., Leinen, M., McManus, J., Pope, R., Smith, C. and Stephens, M.,
1997. Biogenic budgets of particle rain, benthic remineralization and
sediment accumulation in the equatorial Pacific. Deep-Sea Res. Part II, 44:
2251-2281.
Berger, W.H., Adelseck, C.G. and Mayer, L.A., 1976. Distribution of carbonate in
surface sediments of the Pacific Ocean. J. Geophys. Res., 81(15): 26172627.
Boltovskoy, D., 1991. Holocene-Upper Pleistocene radiolarian biogeography and
paleoecology of the equatorial Pacific. Palaeogeog., Palaeoclim.
Palaeoecol., 86: 227-241.
Broecker, W., 2008. Excess sediment 230Th: Transport along the sea floor or
enhanced water column scavenging? Global Biogeochem. Cycles, 22:
Paper GB1006, doi:1010.1029/2007GB003057.
Caress, D.W. and Chayes, D.N., 1996. Improved processing of Hydrosweep DS
multibeam data on the R/V Ewing. Mar. Geophys. Res., 18: 631-650.
Chugh, R.S. and Bhattacharji, J.C., Study of isostasy in the Himalayan Region.
Cochonat, P., Ollier, G. and Michel, J.L., 1989. Evidence for slope instability and
current-induced sediment transport, the RMS Titanic wreck search area,
Newfoundland Rise. Geo-Mar. Lett., 9: 145-152.
Davis, R.E., 2005. Intermediate-depth circulation of the Indian and South Pacific
33
Oceans measured by autonomous floats. J. Phys. Ocean., 35: 683-707.
de Moustier, C., 2001. Field Evaluation of Sounding Accuracy in Deep Water
Multibeam Swath Bathymetry, IEEE Oceans Conference and Exhibition,
Honolulu, Hawaii, USA, Vol. 3, p.1761–1765. OTC 1987, pp. 1761–1765.
de Moustier, C. and Kleinrock, M.C., 1986. Bathymetric artifacts in Sea Beam data:
How to recognize them and what causes them. J. Geophys. Res., 91: 34073424.
Demidova, T.A., Kontar, E.A. and Yubko, V.M., 1996. Benthic current dynamics
and some features of manganese nodule location in the ClarionClipperton Province. Oceanography, 36: 94-101.
Dubois, N. and Mitchell, N.C., 2012. Large-scale sediment redistribution on the
equatorial Pacific seafloor. Deep-Sea Res. I., in press.
Dyer, K.R., 1970. Linear erosional furrows in Southampton Water. Nature, 225:
56-58.
Dymond, J. and Lyle, M., 1994. Particle fluxes in the ocean and implications for
sources and preservation of ocean sediments. In: W.W. Hay, J.T. Andrews,
V.R. Baker, J. Dymond, L.R. Kump, A. Lerman, W.R. Martin, M. Meybeck, J.D.
Milliman, D.K. Rea and F.L. Sayles (Editors), National Research Council:
Material fluxes on the surface of the Earth. Natl Acad. Press, Washington,
D.C., pp. 125-143.
Edwards, M.H., Arvidson, R.E. and Guinness, E.A., 1984. Digital image processing
of SeaBeam bathymetric data for structural studies of seamounts near the
East Pacific Rise. J. Geophys. Res., 89: 11108-11116.
Egbert, G.D. and Erofeeva, S.Y., 2002. Efficient inverse modeling of barotropic
ocean tides. J. Atmos. Oceanic Technol., 19: 183-204.
Eittreim, S.L., 1984. Methods and observations in the study of deep-sea
suspended particle matter. Geol. Soc, Lond., Spec. Pub. 15: 71-82.
Expedition_320/321_Scientists, 2010a. Site U1336. In: H. Pä like, M. Lyle, H. Nishi,
I. Raffi, K. Gamage, A. Klaus and t.E. Scientists (Editors), Proc. IODP,
320/321. Integrated Ocean Drilling Program Management International,
Inc., Tokyo.
Expedition_320/321_Scientists, 2010b. Site U1337. In: H. Pä like, M. Lyle, H. Nishi,
I. Raffi, K. Gamage, A. Klaus and t.E. Scientists (Editors), Proc. IODP,
34
320/321. Integrated Ocean Drilling Program Management International,
Inc., Tokyo.
Farrell, J.W. and Prell, W.L., 1991. Pacific CaCO3 preservation and 18O since 4
Ma: Paleoceanographic and paleoclimatic implications. Paleoceanography,
6: 485-498.
Field, M.E., Gardner, J.V. and Prior, D.B., 1999. Geometry and significance of
stacked gullies on the northern California slope. Marine Geology, 154:
271-286.
Flood, R.D., 1983. Classification of sedimentary furrows and a model for furrow
initiation and evolution. Geol. Soc. Am. Bull., 94: 630-639.
Francois, R., Frank, M., Rutgers van der Loeff, M., Bacon, M.P., Geibert, W.,
Kienast, S., Anderson, R.F., Bradtmiller, L., Chase, Z., Henderson, G.,
Marcantonio, F. and Allen, S.E., 2007. Comment on ‘‘Do geochemical
estimates of sediment focusing pass the sediment test in the equatorial
Pacific?’’ by M. Lyle et al. Paleocean., 22: Paper PA1216,
doi:1210.1029/2005PA001235.
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: doi:
10.1029/2003PA000939.
Gardner, W.D., Sullivan, L.G. and Thorndike, E.M., 1984. Long-term photographic,
current and nephelometer observations of manganese nodule
enivironments in the Pacific. Earth Planet. Sci. Lett., 70: 95-109.
Geibert, W. and Usbeck, R., 2004. Adsorption of thorium and protactinium onto
different particle types: Experimental findings. Geochimica Cosmochimica
Acta, 68: 1489-1501.
Goff, J.A., 2010. Global prediction of abyssal hill root‐mean‐square heights from
small‐scale altimetric gravity variability. J. Geophys. Res., 115: Paper
B12104, doi:12110.11029/12010JB007867.
Griffith, E., Calhoun, M., Thomas, E., Avery, K., Erhardt, A., Bralower, T., Lyle, M.,
Olivarez‐Lyle, A. and Paytan, A., 2010. Export productivity and carbonate
accumulation in the Pacific Basin at the transition from a greenhouse to
icehouse climate (late Eocene to early Oligocene). Paleocean., 25: Paper
35
PA3212, doi:3210.1029/2010PA001932.
Gross, T.F. and Williams, A.J., 1991. Characterisation of deep-sea storms. Mar.
Geol., 99: 281-301.
Hall, I.R., McCave, I.N., Shackleton, N.J., Weedon, G.P. and Harris, S.E., 2001.
Intensified deep Pacific inflow and ventilation in Pleistocene glacial times.
Nature, 412: 809-812.
Hammerstad, E., Pohner, F., Parthiot, F. and Bennett, J., 1991. Field Testing of a
new Deep Water Multibeam Echo Sounder, Oceans '91, pp. 743-749.
Henderson, G.M. and Anderson, R.F., 2003. The U-series toolbox for
paleoceanography. Rev. Mineral. Geochem., 52: 493-531.
Hey, R.N., Kleinrock, M.C., Miller, S.D., Atwater, T.M. and Searle, R.C., 1986.
SeaBeam/Deep-Tow Investigation of an Active Oceanic Propagating Rift
System, Galapogos 95.5°W. J. Geophys. Res., 91: 3369-3394.
Higgins, S., Broecker, W.S., Anderson, R., McCorkle, D.C. and Timothy, D., 1999.
Enhanced sedimentation along the equator in the western Pacific.
Geophys. Res. Lett., 26: 3489-3492.
Hollister, C.D., Flood, R.D., Johnson, D.A., Lonsdale, P. and Southard, J.B., 1974.
Abyssal furrows and hyperbolic echo traces on the Bahama Outer Ridge.
Geology, 2: 395-400.
Honjo, S., Dymond, J., Collier, R. and Manganini, S.J., 1995. Export production of
particles to the interior of the equatorial Pacific Ocean during the 1992
EqPac experiment. Deep-Sea Res., 42: 831-870.
Honjo S., Manganini S.J., Krishfield R.A., Francois R., 2008, Particulate organic carbon
fluxes to the ocean interior and factors controlling the biological pump: A synthesis of
global sediment trap programs since 1983. Progr. Oceanogr. 76:217-285
Izumi, N., 2004. The formation of submarine gullies by turbidity currents. J.
Geophys. Res., 109: Paper C03048, doi:03010.01029/02003JC001898.
Jansen, H., Zeebe, R.E. and Wolf-Gladrow, D.A., 2002. Modeling the dissolution of
settling CaCO3 in the ocean. Global Geochem. Cycles, 16: Paper 1027,
doi:1010.1029/2000GB001279.
Johnson, D.A., 1972a. Eastward flowing bottom currents along the Clipperton
Fracture Zone. Deep-Sea Research, 19: 253-257.
Johnson, D.A., 1972b. Ocean-floor erosion in the equatorial Pacific. Geol. Soc. Am.
36
Bull., 83: 3121-3144.
Key, R.M., Kozyr, A., Sabine, C.L., Lee, K., Wanninkhof, R., Bullister, J.L., Feely, R.A.,
Millero, F.J., Mordy, C. and Peng, T.-H., 2004. A global ocean carbon
climatology: Results from Global Data Analysis Project (GLODAP). Global
Biogeochem. Cycles, 18: paper GB4031, doi:4010.1029/2004GB002247.
Kienast, S.S., Kienast, M., Mix, A.C., Calvert, S.E. and Francois, R., 2007. Thorium230 normalized particle flux and sediment focusing in the Panama Basin
region during the last 30,000 years. Paleocean., 22: Paper PA2213,
doi:2210.1029/2006PA001357.
Kretschmer, S., Geibert, W., Rutgers van der Loeff, M.M. and Mollenhauer, G.,
2010. Grain size effects on 230Thxs inventories in opal-rich and carbonaterich marine sediments. Earth Panet. Sci. Lett., 294: 131-142.
Kunze, E., Firing, E., Hummon, J.M., Chereskin, T.K. and Thurnherr, A.M., 2006.
Global abyssal mixing inferred from lowered ADCP shear and CTD strain
profiles. J. Phys. Ocean., 36: 1553-1576.
Laguros, G.A. and Shipley, T.H., 1989. Quantitative estimates of resedimentation
in the pelagic sequence of the equatorial Pacific. Marine Geol., 89: 269277.
Lonsdale, P., 1976. Abyssal circulation of the southeastern Pacific and some
geological implications. J. Geophys. Res., 81(6): 1163-1176.
Lonsdale, P., 1977. Inflow of bottom water to the Panama Basin. Deep-Sea Res.,
24: 1065-1101.
Lonsdale, P., 1981. Drifts and ponds of reworked pelagic sediment in part of the
Southwest Pacific. Mar. Geol., 43: 153-193.
Lonsdale, P. and Malfait, B., 1974. Abyssal dunes of foraminiferal sand on the
Carnegie Ridge. Geological Society of America Bulletin, 85: 1697-1712.
Lonsdale, P. and Spiess, F.N., 1977. Abyssal bedforms explored with a deeply
towed instrument package. Mar. Geol., 23: 57-75.
Lonsdale, P.F. and Southard, J.B., 1974. Experimental erosion of North Pacific red
clay. Mar. Geol., 17: M25-M34.
Loubere, P. and Richaud, M., 2007. Some reconstruction of glacial-interglacial
calcite flux reconstructions for the eastern equatorial Pacific. Geochem.
Geophys. Geosys., 8: doi:10.1029/2006GC001367.
37
Luo, S. and Ku, T.-L., 2004. On the importance of opal, carbonate, and lithogenic
clays in scavenging and fractionating 230Th, 231Pa and 10Be in the ocean.
Earth Planet. Sci. Lett., 220: 201-211.
Lyle, M., 2003. Neogene carbonate burial in the Pacific Ocean.
Palaeoceanography, 18: DOI: 10.1029/2002PA000777.
Lyle, M., Mitchell, N.C., Pisias, N., Mix, A., Ignacio Martinez, J. and Paytan, A., 2005.
Do geochemical estimates of sediment focusing in the equatorial Pacific
pass the sediment test? Paleoceanography, 20: PA1005,
doi:1010.1029/2004PA001019.
Lyle, M., Pisias, N., Paytan, A., Ignacio Martinez, J. and Mix, A., 2007. Reply to
comment by R. Francois et al. on ‘‘Do geochemical estimates of sediment
focusing pass the sediment test in the equatorial Pacific?’’: Further
explorations of 230Th normalization. Paleocean., 22: Paper PA1217,
doi:1210.1029/2006PA001373.
Lyle, M.W., Owen, R.M. and Leinen, M., 1985. History of hydrothermal
sedimentation at the East Pacific Rise, 19˚S. In: M. Leinen, D.K. Rea and e.
al. (Editors), Init. Repts. DSDP 92. U.S. Govt. Printing Office, Washington,
pp. 585-596.
Macdonald, K.C., Scheirer, D. and Carbotte, S., 1991. Mid-ocean ridges:
discontinuities, segments and giant cracks. Science, 253: 986-994.
Mantyla, A.W., 1975. On the potential temperature in the abyssal Pacific Ocean. J.
Mar. Res., 33: 341-354.
Marcantonio, F., Anderson, R.F., Higgins, S., Stute, M., Schlosser, P. and Kubik, P.,
2001. Sediment focusing in the central equatorial Pacific Ocean.
Paleoceanography, 16: 260-267.
Marcantonio, F., Anderson, R.F., Stute, M., Kumar, N., Schlosser, P. and Mix, A.,
1996. Extraterrestrial 3He as a tracer of marine sediment transport and
accumulation. Nature, 383: 705-707.
Marchal, O., François, R., Stocker, T.F. and Joos, F., 2000. Ocean thermohaline
circulation and sedimentary 231Pa/230Th ratio. Paleocean., 15: 625-641.
Matsumoto, K., 2007. Radiocarbon-based circulation age of the world oceans. J.
Geophys. Res., 112: paper C09004, doi:09010.01029/02007JC004095.
Mayer, L.A., 1980. Deep-sea carbonates: physical property relationships and the
38
origin of high-frequency acoustic reflectors. Mar. Geol., 38: 165-183.
Mayer, L.A., 1981. Erosional troughs in deep-sea carbonates and their
relationship to basement structure. Mar. Geol., 39: 59-80.
Mayer, L.A., Pisias, N.G., Janecek, T.R. and al., e., 1992a. Introduction. In: L.A.
Mayer, N.G. Pisias, T.R. Janecek and e. al. (Editors), Proc. ODP, Init. Repts.
(Ocean Drilling Program), College Station, TX, pp. 5-12.
Mayer, L.A., Pisias, N.G., Janecek, T.R. and al., e., 1992b. Proc. ODP, Init. Repts. 138.
Ocean Drilling Program, Texas A&M University, College Station, TX.
Mayewski, P.A., Meeker, L.D., Whitlow, S., Tickler, M.S., Morrison, M.C.,
Bloomfield, P., Bond, G.C., Alley, R.B., Gow, A.J., Grootes, P.M., Meese, D.A.,
Ram, M., Taylor, K.C. and Wumkes, W., 1994. Changes in atmospheric
circulation and ocean ice cover over the North Atlantic during the last
41,000 years. Science, 263: 1747-1751.
McCave, I.N., 1984. Erosion, transport and deposition of fine-grained marine
sediments. In: D.A.V. Stow and D.J.W. Piper (Editors), Fine-grained
sediments: deep water processes and facies. Blackwell Scientific, Oxford,
pp. 35-69.
McCave, I.N. and Swift, S.A., 1976. A physical model for the rate of deposition of
fine-grained sediments in the deep sea. Geol. Soc. Am. Bull., 87: 541-546.
McGee, D., Marcantonio, F. and Lynch-Stieglitz, J., 2007. Deglacial changes in dust
flux in the eastern equatorial Pacific. Earth Planet. Sci. Letts., 257: 215230.
McGee, D., Marcantonio, F., McManus, J.F. and Winckler, G., 2010. The response of
excess 230Th and extraterrestrial 3He to sediment redistribution at the
Blake Ridge, western North Atlantic. Earth Planet. Sci. Lett., 299: 138-149.
Michaud, F., Chabert, A., Collot, J.-Y., Sallare`s, V., Flueh, E.R., Charvis, P.,
Graindorge, D., Gustcher, M.-A. and Bialas, J., 2005. Fields of multikilometer scale sub-circular depressions in the Carnegie Ridge
sedimentary blanket: Effect of underwater carbonate dissolution? Mar.
Geol., 216: 205-219.
Mitchell, N.C., 1995. Diffusion transport model for pelagic sediments on the MidAtlantic Ridge. J. Geophys. Res., 100(B10): 19,991-920,009.
Mitchell, N.C., 1996. Processing and analysis of Simrad multibeam sonar data.
39
Mar. Geophys. Res., 18: 729-739.
Mitchell, N.C., 1998a. Modeling Cenozoic sedimentation in the central equatorial
Pacific and implications for true polar wander. J. Geophys. Res., 103:
17749-17766.
Mitchell, N.C., 1998b. Sediment accumulation rates from Deep Tow profiler
records and DSDP Leg 70 cores over the Galapagos Spreading Centre. In:
A. Cramp, C.J. MacLeod, S.V. Lee and E.J.W. Jones (Editors), Geological
Evolution of Ocean Basins: Results From the Ocean Drilling Program, Geol
Soc spec publ. Geological Society, London, pp. 199-209.
Mitchell, N.C., Allerton, S. and Escartin, J., 1998. Sedimentation on young ocean
floor at the Mid-Atlantic Ridge, 29˚N. Mar. Geol., 148: 1-8.
Mitchell, N.C. and Huthnance, J.M., 2007. Comparing the smooth, parabolic
shapes of interfluves in continental slopes to predictions of diffusion
transport models. Mar. Geol., 236: 189-208.
Mitchell, N.C., Lyle, M.W., Knappenberger, M.B. and Liberty, L.M., 2003. The
Lower Miocene to Present stratigraphy of the equatorial Pacific sediment
bulge and carbonate dissolution anomalies. Paleoceanography, 18: DOI:
10.1029/2002PA000828.
Moore, T.C., Mitchell, N.C., Lyle, M., Backman, J. and Pälike, H., 2007.
Hydrothermal Pits in the Biogenic Sediments of the Equatorial Pacific
Ocean. Geochem. Geophys. Geosyst., 8: Art. No. Q03015,
doi:03010.01029/02006GC001501.
Morgan, C.L., Allotey Odunton, N. and Jones, A.T., 1999. Synthesis of
environmental impacts of deep seabed mining. Marine Georesources and
Geotechnology, 17: 307-356.
Müller, R.D., Roest, W.R., Roger, J.-Y., Gahagan, L.M. and Sclater, J.G., 1997. Digital
isochrons of the world's ocean floor. J. Geophys. Res., 102: 3211-3214.
Müller, S.A., Joos, F., Edwards, N.R. and Stocker, T.F., 2006. Water mass
distribution and ventilation time scales in a cost-efficient threedimensional ocean model. J. Climate, 19: 5479-5499.
Murray, R.W., Knowlton, C., Leiden, M., Mix, A.C. and Polsky, C.H., 2000. Export
production and terrigenous matter in the Central Equatorial Pacific Ocean
during interglacial oxygen isotop Stage 11. Global Planet. Change, 24: 59-
40
78.
Nozaki Y., Horibe Y., Tsubota H., 1981. The water column distributions of thorium isotopes
in the western North Pacific. Earth and Planetary Science Letters 54:203-216
Nozaki Y., Nakanishi T., 1985. 231Pa and 230Th profiles in the open ocean water column.
Deep-Sea Res. 32:1209-1220
Pälike, H., Lyle, M., Nishi, H., Raffi, I., Gamage, K., Klaus, A. and Shipboard
Scientists, 2010. Proceedings of the Integrated Ocean Drilling Program.
Integrated Ocean Drill. Program, College Station, Tex., pp.
doi:10.2204/iodp.proc.320321.320101.322010.
Pälike, H., Lyle, M.W., Nishi, H., Raffi, I., Ridgwell, A., Gamage, K., Klaus, A., Acton,
G., Anderson, L., Backman, J., Baldauf, J., Beltran, C., Bohaty, S.M., Bown, P.,
Busch, W., Channell, J.E.T., Chun, C.O.J., Delaney, M., Dewangan, P., Jones,
T.D., Edgar, K.M., Evans, H., Fitch, P., Foster, G.L., Gussone, N., Hasegawa,
H., Hathorne, E.C., Hayashi, H., Herrle, J.O., Holbourn, A., Hovan, S., Hyeong,
K., Iijima, K., Ito, T., Kamikuri, S., Kimoto, K., Kuroda, J., Leon-Rodriguez, L.,
Malinverno, A., Moore, T.C., Murphy, B.H., Murphy, D.P., Nakamura, H.,
Ogane, K., Ohneiser, C., Richter, C., Robinson, R., Rohling, E.J., Romero, O.,
Sawada, K., Scher, H., Schneider, L., Sluijs, A., Takata, H., Tian, J., Tsujimoto,
A., Wade, B.S., Westerhold, T., Wilkens, R., Williams, T., Wilson, P.A.,
Yamamoto, Y., Yamamoto, S., Yamazaki, T. and Zeebe, R.E., 2012. A
Cenozoic record of the equatorial Pacific carbonate compensation depth.
Nature, 488: 609-615.
Peine, F., Turnewitsch, R., Mohnc, C., Reichelt, T., Springer, B. and Kaufmann, M.,
2009. The importance of tides for sediment dynamics in the deep sea—
Evidence from the particulate-matter tracer 234Th in deep-sea
environments with differenttidal forcing. Deep-Sea Res. I., 56: 1182-1202.
Puig, P., Palanques, A., Orange, D.L., Lastras, G. and Canals, M., 2008. Dense shelf
water cascades and sedimentary furrow formation in the Cap de Creus
Canyon, northwestern Mediterranean Sea. Cont. Shelf Res., 28: 20172030.
Reid, J.L., 1997. On the total geostropic circulation of the Pacific Ocean: flow
patterns, racers, and transports. Prog. Oceanog., 39: 263-352.
Roussenov, V., Williams, R.G., Follows, M.J. and Key, R.M., 2004. Role of bottom
41
water transport and diapycnic mixing in determining the radiocarbon
distribution in the Pacific. J. Geophs. Res., 109: Paper C06015,
doi:06010.01029/02003JC002188.
Roy-Barman, M., Chen, J.H. and Waserburg, G.J., 1996. 230Th-232Th systematics in
the central Pacific Ocean: The sources and the fates of thorium. Earth
Planet. Sci. Lett., 139: 351-363.
Roy-Barman, M., Jeandel, C., Souhaut, M., Rutgers van der Loeff, M.M., Voege, I.,
Leblond, N. and Freydier, R., 2005. The influence of particle composition
on thorium scavenging in the NE Atlantic Ocean (POMME experiment).
Earth Planet. Sci. Lett., 240: 681-693.
Ryan, W.B.F., Carbotte, S.M., Coplan, J.O., O'Hara, S., Melkonian, A., Arko, R.,
Wiessel, R.A., Ferrini, V., Goodwillie, A., Nitsche, F., Bonczkowski, J. and
Zemsky, R., 2009. Global multi-resolution topography synthesis. Geochem.
Geophys. Geosys., 10: Paper Q03014.
Schmitt, T., Mitchell, N.C. and Ramsay, A.T.S., 2008. Characterizing uncertainties
for quantifying bathymetry change between time-separated multibeam
echo-sounder surveys. Cont. Shelf Res., 28: 1166-1176.
Scott, R.B., Goff, J.A., Naveira Garabato, A.C. and Nurser, A.J.G., 2011. Global rate
and spectral characteristics of internal gravity wave generation by
geostrophic flow over topography. J. Geophys. Res., 116: Paper C09029,
doi:09010.01029/02011JC007005.
Shackleton, N.J., Crowhurst, S., Hagelberg, T., Pisias, N.G. and Schneider, D.A.,
1995. A new late Neogene time scale: Application to leg 138 sites. In: N.G.
Pisias, L.A. Mayer, T.R. Janecek, A. Palmer-Julson and T.H. van Andel
(Editors), Proceedings of the Ocean Drilling Program, Scientific Results,
Vol. 138. Ocean Drilling Program, College Station, TX, pp. 73-101.
Shipley, T.H., Winterer, E.L., Goud, M., Mills, S.J., Metzler, C.V., Paull, C.K. and Shay,
J.T., 1985. Seabeam bathymetric and water-gun seismic reflection surveys
in the equatorial Pacific. In: L. Mayer and F. Theyer (Editors), Init. Repts.
DSDP, 85. U.S. Govt. Printing Office, Washington, pp. 825-837.
Shriver, J.F., Arbic, B.K., Richman, J.G., Ray, R.D., Metzger, E.J., Wallcraft, A.J. and
Timko, P.G., 2012. An evaluation of the barotropic and internal tides in a
high-resolution global ocean circulation model. J. Geophys. Res., 117:
42
Paper C10024, doi:10010.11029/12012JC008170.
Siddall, M., Anderson, R.F., Winckler, G., Henderson, G.M., Bradtmiller, L.I., McGee,
D., Franzese, A., Stoker, T.F. and Müller, S.A., 2008. Modeling the particle
flux effect on distribution of 230Th in the equatorial Pacific. Paleocean., 23:
Paper PA2208, doi:2210.1029/2007PA001556.
Simmons, H.L., 2008. Spectral modification and geographic distribution of the
semi-diurnal internal tide. Ocean Modelling, 21: 126-138.
Singh, A.K., Marcantonio, F. and Lyle, M., 2013. Water column 230Th systematics
in the eastern equatorial Pacific Ocean and implications for sediment
focusing. Earth Panet. Sci. Lett., in press:
doi:10.1016/j.epsl.2012.1012.1006i.
Skinner, L.C. and Shackleton, N.J., 2005. An Atlantic lead over Pacific deep-water
change across Termination I: Implications for the application of the
marine isotope stage stratigraphy. Quat. Sci. Rev., 24: 571-580.
Smith, C.R., Hoover, D.J., Doan, S.E., Pope, R.H., DeMaster, D.J., Dobbs, F.C. and
Altabet, M.A., 1996. Phytodetritus at the abyssal seafloor across 10˚ of
latitude in the central equatorial Pacific. Deep-Sea Res. II, 43: 1309-1338.
Smith, W.H.F. and Sandwell, D.T., 1997. Global sea floor topography from satellite
altimetry and ship soundings. Science, 277: 1956-1962.
Smith, W.H.F. and Wessel, P., 1990. Gridding with continuous curvature splines in
tension. Geophysics, 55(3): 293-305.
Southard, J.B., Young, R.A. and Hollister, C.D., 1971. Experimental erosion of
calcareous ooze. J. Geophys. Res., 76: 5903-5909.
Suman, D.O. and Bacon, M.P., 1989. Variations in Holocene sedimentation in the
North American Basin determined from 230Th measurements. Deep-Sea
Res., 36: 869-878.
Taylor, G.I., 1919. Tidal friction in the Irish Sea. Philosophical Transactions of the
Royal Society, A220: 1-33.
Theyer, F., Vincent, E. and Mayer, L.A., 1989. Sedimentation and
paleoceanography of the central equatorial Pacific. In: E.L. Winterer, D.M.
Hussong and R.W. Decker (Editors), The Eastern Pacific Ocean and
Hawaii. The Geology of North America, v. N. Geological Society of America,
Boulder, Colorado, pp. 347-372.
43
Thomas, E., Turekian, K.K. and Wei, K.Y., 2000. Productivity control of fine
particle transport to equatorial Pacific sediment. Global Biogeochem.
Cycles, 14: 945–955, doi:910.1029/1998GB001102.
Thomsen, L. and Gust, G., 2000. Sediment erosion thresholds and characteristics
of resuspended aggregates on the western European continental margin.
Deep-Sea Res. I., 47: 1881-1897.
Thorndike, E.M., Gerard, R.D., Sullivan, L.G. and Paul, A.Z., 1982. Long-term timelapse photography of the deep ocean floor. In: R.A. Scrutton and M.
Talwani (Editors), The Ocean Floor. John Wiley and Sons, New York, pp.
255-275.
Tominaga, M., Lyle, M. and Mitchell, N.C., 2011. Seismic interpretation of pelagic
sedimentation regimes in the 18–53 Ma eastern equatorial Pacific:
Basin?scale sedimentation and infilling of abyssal valleys. Geochem.
Geophys. Geosys., 12: Paper Q03004, doi:03010.01029/02010GC003347.
Turnewitsch, R., Reyss, J.-L., Nycander, J., Waniek, J.W. and Lampitt, R.S., 2008.
Internal tides and sediment dynamics in the deep sea— Evidence from
radioactive 234Th/238U disequilibria. Deep-Sea Res. I., 55: 1727-1747.
van Andel, T.J., Heath, G.R. and Moore, T.C., 1975. Cenozoic tectonics,
sedimentation, and paleoceanography of the central equatorial Pacific.
Geol. Soc. Am. Mem., 134 pp.
Viekman, B.E., Flood, R.D., Wimbush, M., Faghri, M., Asaka, Y. and Van Leer, J.C.,
1992. Sedimentary furrows and organized flow structure: A study of Lake
Superior. Limnol. Oceanogr., 37: 797-812.
Watkins, J.M., Mix, A.C. and Wilson, J., 1996. Living planktic foraminifera: tracers
of circulation and productivity regimes in the central equatorial Pacific.
Deep-Sea Res. II, 43: 1257-1282.
Watkins, J.M., Mix, A.C. and Wilson, J., 1998. Living planktic foraminifera in the
central tropical Pacific Ocean: articulating the equatorial ‘cold tongue’
during La Nifia, 1992. Mar. Micropal., 33: 157-174.
Wessel, P. and Smith, W.H.F., 1991. Free software helps map and display data.
Eos, Transactions, American Geophysical Union, 72: 441.
Williams, R.G., Day, K., Roussenov, V. and Wood, R., 2002. Role of the overturning
circulation in determining the potential vorticity over the abyssal ocean. J.
44
Geophs. Res., 107: Paper 3170, doi:3110.1029/2001JC001094.
Figures
Figure 1. Maps of the central Pacific, including information on abyssal currents
and sediment properties. (a) Orientations of geomorphological current
indicators are marked by white bars overlying RV Revelle (bold line,
“AMAT03RR”) and other vessel tracks. Background is a map of the bathymetry.
45
White boxes with annotation 2, 3a-3d and 4a-4d locate the data shown in Figures
2-4. Fine lines are tracks where other wide-swath multibeam data have been
collected. Red and purple vectors represent the mean current (scale top-left)
recorded with long-term (>6 month) current meter deployments (>100 mab and
<100 mab, respectively; adjacent values are meter altitudes above bottom). Star
symbols locate the MANOP sites MC, MCD, MS and MM. Site B is a 197-day
current meter measurement from a compilation of Demidova et al. (1996)
reported by Morgan et al. (1999)). Colour-filled circles are values of Y (230Th
“focusing factor”) for the Holocene reported by Kienast et al. (2007). Cross
symbols are locations where tidal currents were calculated for Figure 8.
(b) As (a) with the Last Glacial Maximum values for Y (Kienast et al., 2007)
supplemented with values for the ODP Leg 138 sites 848-853 (McGee et al.,
2007) (marked). Solid vectors show the sense of movement implied by
geomorphologic features. Background is gridded water 14C anomaly (Key et al.,
2004) for 4000 m depth obtained from the Carbon Dioxide Information Analysis
Center (http://cdiac3.ornl.gov/las/servlets/dataset). (Mid-grey region under
"East Pacific Rise" is shallower than 4000 m and contains no data.)
46
47
Figure 2. (a) Shaded relief map of multibeam echo-sounder swath bathymetry
showing elongated troughs ("furrows") trending oblique to abyssal hill trends
(dashed line). Contours are drawn every 50 m with 100 m in bold (depths 3800
and 3900 m are marked). Bathymetry grid has been further filtered with a
Gaussian-weighted filter of full width 0.005˚. Sun illumination is from N055˚E.
Bathymetry cross-sections A-A' etc are drawn at 8.75:1 vertical exaggeration.
See electronic supplement for 10-m contoured version of this figure. (b)
Knudsen chirp profile collected along the swath centre marked EE-E and E-E' in
(a).
48
Figure 3. Shaded relief bathymetry showing examples of furrows (located in
Figure 1a). Arrows mark features discussed in the text. Sun illumination
directions are from (a) N300˚E, (b) N330˚E, (c)) N300˚E, and (d)) N300˚E.
Dashed lines represent interpreted local orientation of abyssal hills. Curves
marked "A" in (c) are centre-beam artifacts where the vessel was turning. See
electronic supplement for contoured versions of these figures.
49
Figure 4. Shaded-relief maps of bathymetry data from on and just south of the
equator (as Figure 2). Aside from some isolated current-generated features
(such as those marked with arrows here), the data on and south of the equator
show less geomorphologic evidence of currents. Sun illumination is from N325˚E
in all these maps. Annotation "P" marks some sedimentary pits. See electronic
supplement for contoured versions of these figures.
50
Figure 5. Selection of current meter records corresponding to the sites marked
MS, (a,d), MC (b,e) and MCD (c) in Figure 1a (“mab” represents metres above
bottom). North current is up on the page. The average current is shown by a
white bar in (a), (b), (d) and (e). Measurement durations and other properties
are given in Table 1. (f) Current speed variation with hours since deployment at
start of the MANOP-C1 20 mab dataset illustrating effect of tidal cycles (vertical
grid shown every 12.42 hours). This is a relatively quiescent part of the dataset,
which elsewhere has a maximum speed of 24.66 cm s-1. The mean displacement
over these 200 hours was to N358.7˚E (almost northward) and progressive
vector diagram shows an initial southward movement followed by a greater
northward movement.
51
52
Figure 6. Data from the MANOP-C sites. (a) Location of sites is shown on
bathymetry (Ryan et al., 2009) contoured every 100 m with background shading
from N030˚E (4000 m in bold). The map contains data from an earliergeneration multibeam system, which did not have the precision needed here to
detect furrows. (b) Temperature records from the two deployments at C1 versus
week from start. (c) Progressive vector diagrams (PVDs) computed from the
current meter data. Solid circles plotted every week from the start, small open
circles every 10 weeks (numbered in places to aid comparisons between
records). PVD in uppermost-right is the C3 42 mab record after removing a
uniform drift representing the annual mean movement. Data provided by the
Buoy Group of the Oregon State University and originally collected with Aandraa
RCM5 meters by J. Dymond as principal investigator. The 42 mab altitude for
MANOP-C3 was estimated from instrument depth and bathymetry in (a) due to
errors in file headers, hence this altitude may be inaccurate.
53
Figure 7. Current meter data from the MANOP-S sites, as Figure 6. (a)
Progressive vector diagrams computed from the current meter data. Data
provided by the Buoy Group of the Oregon State University and originally
54
collected with Aandraa RCM5 meters by J. Dymond as principal investigator. (b)
Locations of sites on coarse bathymetry (Ryan et al., 2009).
Figure 8. Scatter plots of one month of tidal currents calculated from the Oregon
State University TPX06.2 global tide model (Egbert and Erofeeva, 2002) using
MatlabTM Tide Model Driver software at the locations marked by cross symbols
in Figure 1 and at the MANOP C and CD current meter sites.
55
Figure 9. (a) Geochemical 230Th deposition anomalies (focusing factors, Y)
versus latitude for the equatorial region (109˚W to 140˚W) (Kienast et al., 2007;
McGee et al., 2007; Siddall et al., 2008). Solid circles are Holocene values and
open circles are LGM values. Bars on the left are representative uncertainties
derived from data of (unfilled circles) Kienast et al (2007) for their LGM values
and of (solid circles) McGee et al. (2007) for their Holocene values (1). Large
grey-filled and open circles connected by solid lines are averages of Y-values
(Holocene and LGM, respectively) over 1˚ of latitude centred on 0˚ and 1˚N.
Smaller circles connected by dashed lines are similar averages after removing
values highlighted by cross symbols (less reliable values due to potential local
sediment transport). (b) Linear sedimentation rate averages for ten 100-k.y.
intervals of the ODP Leg 138 line of drill sites at 110˚W obtained from the
orbitally tuned timescale of Shackleton et al. (1995). Values have been
56
normalized to the peak value of Site 849 just north of the equator. Bold line
represents the period 0-100 ka. (c) Similar graph to (b) though with data
multiplied by dry bulk density data from the sites (Mayer et al., 1992b) to
convert to mass accumulation rates.
Figure 10. (a) Variation in averaged focusing factors across the equator (open
circles) after removing sites potentially affected by local sedimentary transport.
Dashed line shows averaged factors after reducing three values to illustrate
effect of data uncertainty (see main text).
Grey line is an estimate of the
variation in Y produced by the particle flux effect (Siddall et al., 2008, their
"control" model). (b) Standing stock counts of live foraminifera in net trawls in
57
the central equatorial Pacific around 140˚W (Watkins et al., 1996; Watkins et al.,
1998) (http://www1.whoi.edu/eqpacobjects.html). Abundances were averaged
over the depth interval 0-100 m for February-March 1992 (grey star symbols)
and August-September 1992 (solid star symbols). (c) Carbonate fluxes (plus
symbols) derived from sediment traps between 126˚W and 140˚W (Dymond and
Lyle, 1994; Honjo et al., 1995). (d) Averaged bathymetry (Mitchell, 1998a) for
the region 145˚W-105˚W. (e) Models of 230Th southward flux obtained from the
core Y-values in (a) by solving equation (3) numerically with dy of 1.1 km. Solid
line represents the whole-water column flux (from isopycnal diffusion, water
advection and particle transport). Grey line represents the flux implied by the
particle flux effect (PFE) modelled by Siddall et al. (2008), i.e., grey line in (a).
Dashed line is the predicted meridional sediment-associated 230Th flux predicted
by removing the PFE of Siddall et al. (2008).
Dotted line shows a model
illustrating the effect of data uncertainty (dotted line in (a)).
Figure 11a:
Figure 11b:
58
Figure 11. Total (dissolved and particulate) 230Th activity of water samples.
Measurements were made available by R Anderson, Lamont-Doherty Earth
Observatory at http://usjgofs.whoi.edu. Samples were collected along a
meridional line at 140˚W. (a) Values versus water depth. Grey ellipse highlights
elevated values near the bed discussed in the text. (b) Values were interpolated
to the depths shown. Analytical uncertainties are smaller than symbol size.
Upper graph shows estimates of vertical particle flux corrected to 2000 m depth
from values compiled by Honjo et al. (2008) for the region 135˚W to 140˚W.
Solid circles, open circles and cross symbols represent organic carbon, calcium
carbonate and silica fluxes, respectively.
59
Figure 12. Water 230Th activities from Figure 11a with regression lines through
data excluding lowermost samples. Arrows with associated values highlight the
activity difference of deepest samples from the regression lines.