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Rain ratio variation in the Tropical Ocean: Tests with surface
sediments in the eastern equatorial Pacific$
Figen Mekika,, Paul Loubereb, Mathieu Richaudc
a
Department of Geology, Grand Valley State University, Allendale, MI 49401, USA
Department of Geology and Environmental Geosciences, Northern Illinois University, DeKalb, IL 60115, USA
c
Department of Geology and Geography, Georgia Southern University, 1110 Herty Building, Herty Drive, Statesboro, GA 30460-8149, USA
b
Accepted 11 January 2007
Abstract
The organic carbon to calcite flux ratio (rain ratio) has a profound effect on the preservation of carbonates in the deep
sea and may influence atmospheric pCO2 over millennia. Unfortunately, the degree to which the rain ratio varies in the
more productive regions of the oceans is not well determined with sediment trap data. The rain ratio in the upper ocean
appears dominantly linked to diatom productivity, which is not necessarily directly linked to total production and may be
regionally variable. However, ballasting and protection of organic carbon by calcareous particles in the deeps may limit
ratio variability at the seafloor. Sediment trap data do not exist for the regional determination of rain ratios in key highly
productive areas like the eastern equatorial Pacific (EEP). To overcome this, we turn to surface sediment composition and
accumulation rates as a representation of modern ratio variation.
We present 230Thorium (230Th)-normalized carbonate, opal, organic carbon and detrital matter accumulation rates from
core top samples in the EEP. We demonstrate a novel approach for estimating modern rain ratios from sedimentary
proxies by (1) calculating vertical calcite flux from 230Th-normalized carbonate accumulation rates (CARs) with correction
for preservation and (2) calculating organic carbon fluxes with multiple algorithms that depend in varying degrees on
ballasting. We find that organic carbon flux estimates from algorithms with and without a ballasting function produce
results different from one another. Sediment accumulation rates for opal reflect the likely pattern of diatom production. By
contrast, the organic carbon accumulation rate does not correlate well with surface ocean productivity or any of our
algorithm-based organic carbon flux estimates. Instead, it correlates with the detrital component of the sediments
suggesting an allochthonous input to sedimentary organic carbon accumulation in the EEP, which reduces its value as a
productivity tracer. However, our calcite and multiple, satellite-based organic carbon fluxes allow estimation of the rain
ratio and demonstrate a common regional pattern with moderate to strong variability in the rain ratio across the EEP. This
variability is significant and is transmitted into the deeps leaving a sedimentary record regardless of the algorithm chosen
to calculate organic carbon fluxes. Furthermore, we provide evidence suggesting that the rain ratio in the EEP may be
driven by wind-supplied iron availability, which would regionally enhance nutrient use and promote diatom growth.
r 2007 Elsevier Ltd. All rights reserved.
Keywords: Rain ratio; Tropical Pacific; Multiproxy
$
Tables for all data and all calculations of organic carbon fluxes and rain ratios are available upon request from the lead author or by
visiting http://www4.gvsu.edu/mekikf.
Corresponding author.
E-mail address: [email protected] (F. Mekik).
0967-0645/$ - see front matter r 2007 Elsevier Ltd. All rights reserved.
doi:10.1016/j.dsr2.2007.01.010
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1. Introduction
The ratio of vertical fluxes of organic carbon to
calcite to the deep sea, known as the rain ratio, is
important for the oceanic regulation of atmospheric
pCO2 over millennia (Archer and Maier-Reimer,
1994). Developing tools for accurately reconstructing the rain ratio in sediments is important because
tracking changes in the rain ratio and the preservation state of deep-sea carbonates in down-core
samples can enable the calculation of bottom water
DCO3¼ ([CO3¼ ]in situ minus [CO3¼ ]saturation); and
this, in turn, may lead to better understanding the
marine carbon cycle component driving atmospheric pCO2 changes over time. Additionally,
tracking changes in the rain ratio through time
would help reveal how marine pelagic communities
respond to climate change. Community changes
might be driven by variation in the chemistry of
equatorial undercurrents (Matsumoto et al., 2002),
source regions for upwelled water (Loubere, 2000,
2001), or supply of aeolian micronutrients (e.g.,
wind-blown iron; Vink and Measures, 2001).
However, the rain ratio in the deep ocean may
remain relatively constant if calcareous particles act
as the dominant ballast for organic carbon through
the twilight zone (Armstrong et al., 2002; Franc- ois
et al., 2002; Klaas and Archer, 2002). This would
mean the rain ratio influence on the deep-ocean
carbonate system could only be modest (Ridgwell,
2003). Also, it would imply that reconstruction of
past variations in upper-ocean ratios and phytoplankton community structure from deep-sea sediments could be difficult due to signal attenuation.
We undertake to test for deep ocean variation in
the rain ratio signal preserved in deep-sea sediments. Data on the shallow and deeper rain ratios
of the more highly productive regions of the oceans
are scarce. This is especially true for the eastern
tropical oceans, which appear to play an important
role in ocean-atmosphere CO2 exchange (Takahashi
et al., 2002) and new nutrient supply to the surface
ocean (Sarmiento et al., 2004). A key region in terms
of the marine carbon cycle (e.g., Toggweiler and
Carson, 1995), biological productivity (Chavez and
Barber, 1987) and phytoplankton community structure (Ragueneau et al., 2000; Wilkerson and Dugdale, 1996) is the eastern equatorial Pacific (EEP).
The EEP is a unique region of the world ocean
where phytoplankton community structure and
productivity are controlled by a complex interaction
of nutrient supply, upwelling dynamics and avail-
ability of key micronutrients, like Fe. An important
source for most nutrients in the EEP seems to be
waters from the SW Antarctic Pacific (Toggweiler
et al., 1991). These nutrients reach the surface
through deep upwelling off Peru, and then advect
westwards in the South Equatorial Current. However, all of these nutrients are not used up by the
biota. This characteristic makes the EEP a highnutrient low-chlorophyll (HNLC) region. Windsupplied Fe may strongly affect the distribution of
diatom vs. coccolithophore waters in this region
because Fe is a limiting micronutrient (Martin and
Fitzwater, 1988; Coale et al., 1996; Martin et al.,
1994; Fitzwater et al., 1996).
It is important to determine the variability of rain
ratios in upwelling regions like the EEP because (1)
there is a strong gradient in surface-ocean productivity there along with the HNLC condition, which
leads to high-diatom productivity on the Peruvian
margin and restriction on diatoms in the open ocean
and (2) there is a strong efflux of CO2 (Tans et al.,
1990; Takahashi et al., 2002) into the atmosphere
driven by upwelling. Unfortunately, there is currently no regionally extensive sediment trap dataset
for the EEP. To substitute for the lack of watercolumn data, we present a novel approach for
reconstructing the rain ratio using sedimentary
proxies and we apply our approach to core top
(modern) samples from the EEP.
Reconstruction of the rain ratio at the seabed
from sedimentary proxies is a challenge. To estimate
the ratio, we need to calculate the original fluxes of
both labile organic carbon and calcite. We approach
the former by using satellite-derived productivity
estimates and published algorithms for flux vs.
water depth. To estimate calcite flux, we turn to the
accumulation rate of calcite in surface sediments.
However, there are three complicating factors: (1)
calculating the percent of calcite dissolved in the
sediment [calcite flux ¼ accumulation rate/fraction
of calcite preserved], (2) correcting for lateral
sediment redistribution; so-called focusing, which
biases estimates of sediment accumulation rate due
to direct ‘‘overhead’’ supply (Franc- ois et al., 2004),
and (3) calculating the degradation of organic
carbon through the water column and therefore
the flux of organic carbon from the surface ocean to
the seabed. We address the first two issues by using
a recently developed carbonate preservation proxy,
the Globorotalia menardii fragmentation index (after
Mekik et al., 2002) and 230Thorium (230Th)-normalized carbonate accumulation rates (CARs) in order
Please cite this article as: Mekik, F., et al., Rain ratio variation in the Tropical Ocean: Tests with surface sediments in the eastern
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to correct for both carbonate preservation and
sediment redistribution at any point on the seafloor
(Franc- ois et al., 2004).
The third factor is more complicated because the
nature and magnitude of carbon degradation both
in the water column and in sediments are difficult to
assess. Many algorithms have been published for
calculating water-column organic carbon flux from
surface-ocean productivity estimates (the ones we
test here are Berger et al., 1988, 1989; Antia et al.,
2001; Franc- ois et al., 2002; Klaas and Archer,
2002). All of these algorithms are derived from
sediment trap data at various water depths. Francois et al. (2002) and Klaas and Archer (2002),
following the work of Armstrong et al. (2002),
tested the effect of mineral ballast in facilitating
organic carbon transport to the seabed. Although
the Franc- ois et al. (2002) algorithm takes f-ratios as
well as mineral ballast into account, both Franc- ois
et al. (2002) and the Klaas and Archer (2002) found
calcite particles to be important ballasting material
for organic carbon. We should note that unlike
Franc- ois et al. (2002), in Klaas and Archer (2002)
algorithm, detrital sediment fluxes play an equally
significant role for ballast. Nonetheless, if calcareous particles function as ballast and protection for
organic carbon to the seafloor, then this would
significantly limit the variability of the rain ratio at
the seabed. Conversely, algorithms like those of
Berger et al. (1988, 1989) and Antia et al. (2001) do
not assume any ballasting mechanism and are based
solely on primary production and water depth.
The questions we examine in the EEP with our
multi-proxy approach are:
(1) Is there consistent evidence for regional variation in the rain ratio for the EEP?
(2) Does the rain ratio reconstructed with sedimentary proxies strictly follow trends in organic
carbon fluxes?; and if not, what is revealed
about phytoplankton community structure in
the EEP?
2. Materials and methods
2.1. Samples
Fig. 1 illustrates the distribution of our samples in
the EEP. All samples are from gravity cores and
taken from the top 0–4 cm of each core. We chose
gravity cores exclusively to limit the possibility of
3
Fig. 1. Geographic distribution of surface sediment samples used
herein. Gray and black squares are samples with 230Th data,
black dots are those with % opal estimates. White squares are
samples in addition to the others for which there are G. menardii
fragmentation data.
sediment loss while coring. We excluded two
samples from our batch (not shown in Fig. 1)
because their d18O values were within the range of
values known for sediments from the last glacial
maximum (LGM). Our sample set for carbonate
preservation, % calcite and % organic carbon in dry
bulk sediment contains 50 samples. We generated
230
Th data for a subset of these samples (37; gray
and black samples in Fig. 1) and % opal data for a
subset of those (20; black dots in Fig. 1).
2.2. Data
All % calcite, % organic carbon, and % opal
data were produced at Northern Illinois University.
Percent calcite and organic carbon were measured
on the Carlo-Erba NA 1500 C/N/S following the
procedure of Verardo et al. (1990). Samples were
calibrated against standard acetanilide and checked
against NBS standard reference material 1b (argillaceous limestone). Replicate analyses of the sediments indicate a mean error of 72%. Percent opal
was measured using the method of Mortlock and
Froelich (1989) and single extraction of silica into a
Na2CO3 solution. Every sample was analyzed twice
(three times when results diverge) and the uncertainty is about 74%.
Percent preserved carbonate was calculated using
the G. menardii fragmentation index (MFI after
Mekik et al., 2002) which is a dissalution proxy with
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a strong quantitative relationship calculated with
decreasing DCO3¼ [R2 ¼ 0.89] in its calibration
areas on the Ontong–Java Plateau and the East
Pacific Rise. MFI also was anchored against
modeled estimates of percent carbonate dissolved
using the biogeochemical model, Muds (Archer
et al., 2002) (R2 ¼ 0.88) (see Mekik et al., 2002 for
details of modeling). MFI was calibrated in areas
outside the upwelling zone where the dominant
dissolution driver is DCO3¼ . Subsequently, Mekik
et al., (2002) applied MFI to a 46-sample set from
the EEP and modeled the rain ratio. In their
modeled rain ratio map, the imprints of the South
Equatorial Current and the North Equatorial
Counter-current are clearly discernable. However,
they did not have 230Th-normalized sediment
accumulation rate and vertical calcite flux data at
that time.
MFI ¼ D=ðD þ W Þ,
(1)
where D is the number of damaged G. menardii
fragments and W is the number of whole G.
menardii specimens
D ¼ # greater than half þ ½# less than half=3
þ ½# of keels=5.
ð2Þ
The MFI transfer function equation is
% calcite dissolved ¼ 5:111 þ ½MFI 160:491
½MFI2 79:636.
ð3Þ
The precision of MFI is 75.8% dissolved, and its
accuracy is 710%. Our MFI data are based on 300
or more counts of fragments and whole tests per
sample for the majority of our samples. Few
samples have fewer G. menardii tests and few
Table 1
Equations for
samples with very high dissolution only contained
keels and very few whole specimens. In those
samples counts fell below 300 due to lack of
material.
230
Th data were generated at Woods Hole
Oceanographic Institution using the isotope dilution
method of Choi et al. (2001) and measured on a
thermofinnigan element ICP-MS. Samples were
prepared by acid digestion of sediment and thorium
separation by anion exchange (Anderson and Fleer,
1982). Precision of the measurements is better than
2%.
Table 1 lists the relationships and equations we
use herein for calculating 230Th-normalized sediment accumulation rates and vertical fluxes.
We used satellite-based data from Behrenfeld and
Falkowski (1997) to estimate primary productivity
(PP) for each of our sample locations. We used a
multiple algorithm approach (as outlined in Table
2) for converting PP to seabed organic carbon flux.
The f-ratios used in Franc- ois et al.’s (2002)
algorithm for each of our sampling locations were
kindly provided by Richard Krishfield (personal
communication, 2006) based on the algorithms in
Laws et al. (2000). We used these f-ratios to convert
PP values to export production estimates EP.
Calcite flux estimates in both Franc- ois et al.
(2002) and Klaas and Archer (2002) algorithms
were calculated by correcting 230Th-normalized
CAR with carbonate preservation as denoted in
Table 1. Methods for estimating % opal preserved
in sediments is currently at its infancy (Pichon et al.,
1992; Sayles et al., 2001; Dezileau et al., 2003)
because opal dissolves both in the water column and
within sediments (Reed Scherer, personal communication, 2006). Furthermore, early studies on opal
230
Thorium-normalized sediment accumulation rates and vertical fluxes
Parameter of interest
Equation
Explanation of terms
Bulk sedimentation rate [BSR]
BSR ¼ (b* water depth [km])/ex230Tho
b ¼ constant production rate of 230Th from 234U
b ¼ 2.63 dpm/cm2/ka/km of water depth
ex230Tho ¼ original 230Th activity in sediments in g/
dpm.
Carbonate accumulation rate [CAR]
Opal accumulation rate [OAR]
Organic C accumulation rate [Org
CAR]
Detrital sediment accumulation rate
[DSAR]
Vertical calcite flux [VCF]
CAR ¼ BSR*fraction carbonate
OAR ¼ BSR*fraction opal
Org CAR ¼ BSR*fraction organic
carbon
DSAR ¼ BSR–CAR–OAR–Org CAR
% sediment component ¼ % by dry weight
VCF ¼ CAR/fraction calcite
preserved
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Table 2
Comparison of four algorithms for calculating organic carbon flux to the seabed. Equations are represented with the same notation used in
their respective manuscripts
Algorithm
Data base
Assumptions
Berger et al., 1988
J(z) ¼ 0.17 (PP/(z/10))+0.01PP, where
PP ¼ primary production z ¼ water depth in
meters
Global sediment trap data
below 1000 m
f-ratio is constant or has low variability
Org. C flux to the deeps is controlled by:
(1) export production and
(2) water depth
No ballasting component
Sediment traps from 99 time
series in 27 sites
All Atlantic data from 801N to
651S
Mostly from high surface-ocean
productivity regions
Variability in f-ratio
Global sediment trap data
below 2000 m
57 locations
Variability in f-ratio
Antia et al., 2001
J (Corg) ¼ (0.1 PP1.77) (z0.68), where
PP ¼ primary production z ¼ water depth in
meters
Franc- ois et al., 2002
Teff ¼ 2.51 103 Fcarb+102/z0.096 fratio+0.009,
where z ¼ water depth in meters Teff ¼ transfer
efficiency Fcarb ¼ vertical calcite flux
FCorg ¼ Teff*EP,
where EP ¼ export production
Klaas and Archer, 2002
FC ¼ (0.025 opal flux)+(0.074 calcite
flux)+(0.071 clay flux)+b (PP/z),
where PP ¼ primary production
z ¼ water depth in meters
b ¼ proportionality coefficient of POC to PP
b is negligible below 2000 m
Global long-term sediment trap
data
Both as time series and annual
flux data
52 locations
1000–4833 m water depth
dissolution based on diatom assemblages from the
Southern Ocean are difficult to apply to sediments
from the EEP. Thus, we could not convert opal
accumulation rates to vertical opal fluxes, which
limits the accuracy of our results from the Klaas and
Archer (2002) algorithm. However, the coefficient
for opal flux in their algorithm is significantly
smaller (bya third) than that for calcite or detrital
matter (Table 2). Francois et al. (2002) found opal
flux to be negligible as a ballasting component in
their algorithm. Detrital sediment accumulation
rates for our samples were determined by subtracting 230Th-normalized accumulation rates for biogenic sedimentary components in our samples from
our bulk 230Th-normalized sediment accumulation
rate estimates. Because detrital matter does not
dissolve in sediments, detrital matter accumulation
rate ¼ vertical flux of detrital matter. The exact
Org. C flux to the deeps is controlled by:
(1) export production and
(2) water depth
No ballasting component
Org. C flux to the deeps is controlled by
(1) export production and
(2) water depth
Calcite particles act as ballast for org C
Extension of Armstrong et al.’s (2002) model
Distinguishes different forms of mineral ballast
Dominantly calcite and clay
equations we used for each algorithm are listed in
Table 2.
2.3. Modeling
The rain ratio has a significant effect on the
preservation of calcite in deep-ocean sediments (e.g.,
Berger, 1992; Archer, 1996a; Mekik et al., 2002).
Generally, all other factors being equal, increasing
the ratio leads to increasing dissolution within the
sediments. Hence, we can use the estimated calcite
% dissolved from the surface sediment MFI proxy
to test for the potential variation in the rain ratio
across the EEP. This is done by modeling the
expected % calcite dissolved for our samples with a
constant rain ratio. Then, we compare this to our
observed % dissolved to see if there are significant
differences between the two with a regional pattern
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that we might expect for rain ratio changes
associated with EEP productivity gradients. The
differences between model results and observations
can be examined in terms of the types of rain ratio
changes needed to explain them.
We used Archer et al. (2002) biogeochemical
model, muds, for this exercise. Muds is a sediment
redox diagenesis model that allows the integration
of sedimentary properties measurable in cores with
biogenic fluxes and bottom-water properties important to our understanding of the marine carbon
cycle. The model and its development are fully
explained in Archer et al. (2002). The model applies
to slope and abyssal sediments of the Pacific and
Atlantic. It uses steady-state diffusion-reaction
equations and previously determined reaction kinetics. It includes oxic, sub-oxic and anoxic
respiration. The model uses first-order dissolution
kinetics for calcite. Where there is uncertainty, the
model parameterizes rate constants as functions of
sediment respiration rate. These parameterizations
are tuned to best fit a 53-site calibration data set.
The calibration sites include the Manop locations in
the tropical Pacific. For our purposes, we can use
Muds to calculate calcite dissolution rates for the
organic carbon fluxes and bottom-water DCO3¼
values of our samples. We used a modification of
Muds [Muds_constcal], where the input parameters,
for each sample location are water depth, organic
carbon flux, % calcite in sediments and DCO3¼ ; and
the output parameter is calcite dissolution rate. We
obtain organic carbon fluxes from the algorithms in
Table 2, as explained above, and we use DCO3¼
values from Archer (1996a, and personal communication, 2001). We can estimate the fraction of
calcite dissolved by: % calcite dissolved ¼ dissolution rate/ vertical calcite flux. We calculate the
vertical calcite flux as a constant multiple of the
organic carbon flux as we are assuming a constant
rain ratio for the modeling exercise. Following
Klaas and Archer (2002) and Mekik et al. (2002),
we use a rain ratio of 0.6 [so calcite flux ¼ 1.67 organic carbon flux].
although topography is variable across the region.
Calcite dissolution here will be driven by both
spatial variations in the rain ratio and bottom water
DCO3¼ . In regions outside the upwelling zone, we
see a range of variability for % calcite preserved
from 25% to 70% both in our samples herein and in
Mekik et al.’s (2002) sample set.
3.2. 230Th-normalized sediment accumulations rates
in the EEP
Fig. 2 shows 230Th-normalized accumulation
rates in the EEP for three sediment components.
All our maps are superposed onto the surface-ocean
productivity map (standard model) from Behrenfeld
and Falkowski (1997). CAR (Fig. 2(A)) shows a
distinct latitudinal pattern that will be a function of
original calcite flux, rain ratio and water depth. The
latter two control carbonate preservation. The opal
accumulation rate pattern (Fig. 2(B)) follows zones
of high-surface-ocean productivity. This also is a
function both of original flux and preservation. The
latter is complex (Archer et al., 1993) and we do not
address it here. Organic carbon accumulation rates
calculated from % organic carbon in dry bulk
sediment and 230Th-normalization are shown in
Fig. 2(C). Vertical labile organic carbon fluxes are
discussed below.
3.3. 230Th-normalized vertical calcite flux to the
seabed in the EEP
By dividing the 230Th-normalized CARs
(Fig. 2(A)) with MFI-based estimates of percent
calcite preserved, we generate a vertical calcite flux
map for the EEP (Fig. 2(D)). Vertical calcite fluxes
in the EEP not only follow a latitudinal pattern but
also show strong meridional gradients nearer to
South America. Calcite fluxes are high along the
equator and drop northward and southward away
from the equator. This north–south gradient to
calcite flux also exists along Peru where fluxes are
high near the equator and north of it, but low south
of the equator.
3. Results and discussion
3.4. Organic carbon flux to the seabed in the EEP
3.1. Carbonate preservation in the EEP
Percent calcite preserved is calculated with MFI
in our samples. Note that high-upwelling regions of
our study area have a more or less homogeneous
carbonate preservation pattern (25–30% preserved)
We used PP estimates from Behrenfeld and
Falkowski (1997) to estimate organic carbon fluxes
in the EEP. Table 3 lists the correlation matrix
for these different organic carbon flux estimates.
The estimates fall into two groups having low
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Fig. 2. 230Th-normalized sediment component accumulation rates and reconstructed calcite flux. Accumulation rates in g/cm2 kyr: (A)
calcite accumulation rate, (B) opal accumulation rate, (C) organic carbon accumulation rate and (D) 230Th-normalized vertical calcite flux
in mmol/cm2 yr, units selected to be compatible with organic carbon flux estimates and calculation of the rain ratio (mmol/mmol). Base map
is that of productivity from the standard model of Behrenfeld and Falkowski (1997).
Table 3
Correlation matrix for organic carbon flux calculated with various algorithms
Correlations
Berger et al., 1988,
1989
Antia et al., 2001
Franc- ois et al., 2002
Klaas and Archer,
2002
Berger et al., 1988, 1989
Antia et al., 2001
Franc- ois et al., 2002
Klaas and Archer, 2002
1.00
0.98
0.48
0.31
0.98
1.00
0.41
0.32
0.48
0.41
1.00
0.73
0.31
0.32
0.73
1.00
correlations to one another: Berger et al. (1988,
1989) and Antia et al. (2001) fall into the nonballasting group; while Franc- ois et al. (2002) and
Klaas and Archer (2002) algorithms have strong
ballasting components. Part of the lower correlation
between these two groups may be a product of the
different number of samples for which calculations
could be made (lower number of samples for
ballasting models). We cannot calculate organic
carbon fluxes using Franc- ois et al. (2002) algorithm
for samples in the Peru upwelling region because we
do not have f-ratio estimates for those samples.
Sampling for the Klaas and Archer (2002) algorithm
is limited to those for which % opal data are
available. We use one algorithm representative of
each group to reconstruct rain ratios in the EEP.
These are those of Berger et al. (1988, 1989) and
Franc- ois et al. (2002).
Note that the underlying assumptions of the
Berger et al. (1988, 1989) and Franc- ois et al. (2002)
algorithms are different; and their organic carbon
flux patterns across the EEP follow different trends
(Fig. 3). Franc- ois et al. (2002) algorithm produces
lower organic carbon fluxes than those of the Berger
et al. (1988, 1989) algorithm in higher productivity
regions (Fig. 3(C)). The export production calculated by Berger et al. (1988, 1989) which is simply
10% of primary production, and that calculated
with the Franc- ois et al. (2002) algorithm using
f-ratios based on Laws et al. (2000) algorithms and
sea-surface temperature are strongly correlated
(Fig. 3(D)), so differences between the two methods
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Fig. 3. Organic carbon fluxes to the seabed calculated using algorithms in Table 1. Units in mmol/cm2 yr: (A) flux based on Berger et al.
(1988, 1989), (B) flux based on Franc- ois et al. (2002), (C) differences in the fluxes of A and B [A–B] and (D) export flux comparison
between the algorithms of Berger et al. (1988, 1989) and Franc- ois et al. (2002).
in the seabed organic carbon fluxes are the result of
ballasting effects in the Franc- ois et al. (2002)
calculations.
Fig. 4 shows plots of vertical calcite flux against
organic carbon flux using both algorithms. Note
that the organic carbon fluxes from Franc- ois et al.
(2002) algorithm have a significant linear relationship with vertical calcite flux, and organic carbon
fluxes with Berger et al. (1988, 1989) algorithm have
no such relationship. This is the result of the calcite
ballast component in Franc- ois et al. (2002) equations linking the organic carbon and calcite fluxes.
3.5. Sedimentary organic carbon accumulation rate
as a tracer of organic carbon flux
To examine the rain ratio, and its potential
changes through time, we need estimators for both
original calcite and labile organic carbon fluxes. We
seek tracers of the vertical flux through the water
column, representing overlying biological productivity and for organic carbon we seek the labile,
reactive component. This is the portion of the
organic carbon flux that will play a part in the
carbon cycling of the deep ocean. It has generally
been assumed that organic carbon accumulation
rates serve as a proxy for the vertical water-column
flux of reactive organic matter. In Fig. 5, we
compare the organic carbon accumulation rates
we have measured with the water-column organic
carbon fluxes estimated with the Berger et al. (1988,
1989) and Franc- ois et al. (2002) algorithms. The
relationship is poor in each case, suggesting that
the relationship between organic carbon flux to the
seabed and accumulation rates is not a simple one in
the EEP. Fig. 2 shows that regional organic carbon
accumulation rates show a strong east–west gradient, with higher values extending westward in
regions that have relatively lower surface-ocean
productivities (Panama Basin; south of the equator). There is no ‘tongue’ of higher values tracking
the band of higher productivity along the equator.
The regional pattern of organic carbon accumulation rates does not resemble any map of primary
production that has been made for the EEP based
on field observations (summarized in Berger et al.,
1989) or satellite data (Behrenfeld and Falkowski,
1997; Antoine et al., 1996).
The possible explanations for this are: (a) the
regional pattern for organic carbon export in the
EEP differs substantially from that of primary
production, (b) preservation factors complicate the
organic carbon flux and accumulation rate connection, or (c) there is allochtonous input of organic
carbon so that ‘overhead’ supply is not the sole
source of organic carbon to the sediments.
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9
Fig. 4. Comparison of organic carbon fluxes from the (A) Berger et al. (1988, 1989) and (B) Franc- ois et al. (2002) algorithms and the
reconstructed calcite flux for the EEP.
For export flux to differ greatly from primary
production in the EEP, and for this difference to
account for sediment organic carbon accumulation
rates (Fig. 2(C)), export would have to be relatively
high off the equator and become systematically
lower for the equatorial higher-production belt.
There is no obvious reason for this to be the case,
and analysis of export ratios for the EEP does not
suggest it (Laws et al., 2000). If such a pattern of
export change across the equator did exist, it might
be driven by a ballasting process. If so, we would
expect to find a relationship between calcite
accumulation rates and fluxes and organic carbon
accumulation rates, since calcite appears to be the
dominant ballast material. We find no such
correlation among these variables. Nor do we find
any useful correlations between sediment/calcite
accumulation rates or fluxes and the residuals for
regression of organic carbon accumulation rates
against calculated organic carbon fluxes (Fig. 5). In
Fig. 2, it can be seen that the regional pattern of
organic carbon accumulation rates is not similar to
the calcite, or the opal data. Finally, analysis of
benthic foraminiferal assemblages in surface sediments of the EEP has demonstrated a strong
compositional response that is well related to the
surface-ocean productivity (Loubere, 1994; Loubere
and Fariduddin, 1999) and in which the changes in
faunal elements follow the expected for a response
to flux of labile organic carbon to the seabed. This
faunal signature demonstrates that the pattern of
export flux in the EEP cannot be significantly
different from that of PP, as seen in ship-based
and satellite studies.
Preservational factors may well influence the
organic carbon accumulation rates we have measured. Jahnke (1996) discussed the regional patterns
in preservation that seem to be imposed by different
sedimentation regimes in the deep sea. There are
issues of oxygen supply, sediment accumulation
rate, degradation rates for organic materials, and
possibly sediment composition. We are not currently in a position to analyze these variables, but
we do not find a simple relationship between
organic carbon accumulation rates and calcite or
sediment accumulation rates, or calcite fluxes.
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Fig. 5. Comparison of the sediment organic carbon accumulation rates with the seabed organic carbon fluxes estimated by the (A) Berger
et al. (1988, 1989) and (B) Franc- ois et al. (2002) algorithms. We find no statistically significant relationship.
Neither do we find a relationship between these
variables and the differences between organic
carbon accumulation rates and estimates seabed
flux. Thus, preservation remains an open question.
Allochthonous sources for organic carbon are
also a possibility. In Fig. 6, we compare the
sediment organic carbon accumulation rate with
the accumulation rate of detrital material (total
sediment accumulation rate minus the sum of the
biogenic sediment accumulation rates). The relationship here is strong and it cannot be explained as
a by-product of the calculation method for the
detrital component because the fraction of the total
accumulation rate accounted for by organic carbon
is small. So, the accumulation rate of the detrital
component is not controlled by this. These results
indicate that a process related to the detrital
component, most likely lateral transport of sediments through the atmosphere and the water
column, is significantly influencing organic carbon
accumulation rates. Rea (1994) examined the dust
record of the tropical Pacific and found evidence
that substantial input via water-column transport
occurred in the EEP. This also may affect organic
carbon as it travels with the finer fraction of the
sediments. Our results indicate that organic carbon
accumulation rates cannot be taken as a simple
proxy for the original water-column organic carbon
flux.
3.6. Using calcite preservation to test for rain ratio
variability
Since the rain ratio impacts calcite preservation at
the seabed, we can use surface sediment data to
detect rain ratio variations. Other factors being
equal, increasing the rain ratio leads to reduced
calcite preservation because organic carbon oxidation in the seabed generates acids which cause
dissolution. We can test for ratio variability by
examining regional trends in calcite preservation,
and comparing these to trends expected if the rain
ratio were constant across the EEP. As explained in
the methods section, we can use modeling to predict
calcite preservation at the seabed given the organic
carbon fluxes derived from the algorithms reviewed
above. Using the model allows us to control for the
influence of bottom water saturation on calcite
dissolution and to identify the effects of the rain
ratio. Calcite preservation predicted with the model
can be compared to that estimated for the seabed
from the MFI proxy. Fig. 7 shows the MFI %
Please cite this article as: Mekik, F., et al., Rain ratio variation in the Tropical Ocean: Tests with surface sediments in the eastern
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Fig. 6. Comparison of
11
230
Th-normalized organic carbon and detrital component accumulation rates for sediments in the EEP.
calcite preserved values minus the model predicted
% preserved for a constant rain ratio, plotted
against organic carbon flux predicted by the Berger
et al. (1988, 1989) and Franc- ois et al. (2002)
algorithms. The results for the two different carbon
flux data sets show a trend with model preservation
becoming increasingly better than preservation
actually observed in the sediments as organic
carbon flux increases. Modeling shows that for a
constant rain ratio [ ¼ 0.6], calcite preservation
should improve as surface-ocean productivity increases. The reason for this is that, for a constant
rain ratio, the supply of calcite to the seabed
outstrips the combined dissolution generated by
bottom-water saturation state and within sediment
organic carbon oxidation. Increased burial rate with
higher sedimentation rates also plays a part in better
calcite preservation. This model prediction is far
from what is observed in more productive areas of
the EEP where calcite preservation is often poor, as
indicated by the MFI proxy and simple visual
observations of the sediments.
Fig. 8 presents a map of the observed to modeled
preservation difference across the EEP. The exact
value of the difference is a function of the rain ratio
used in the model. As explained in the methods
section, we used a value of 0.6 for the rain ratio.
Increasing this value will decrease the modeled %
preserved, which would shift the preservation
difference on Fig. 8 to more positive values. The
reverse would occur if the rain ratio used in the
model were less than 0.6. Changing the rain ratio
will have no effect on the overall trends in
preservation difference between the MFI index
and modeled values. The index to model difference
has a coherent regional pattern, becoming increasingly negative (preservation from the model greater
than that observed in the sediments) towards the
equator and towards the continental margin of
the Americas. The range of this difference exceeds
the uncertainty in the MFI sedimentary index
(about 10%) and is therefore significant. The
discrepancy between modeled and observed values
could be resolved in a simple way by varying the
rain ratio across the EEP and having higher values
towards zones of higher productivity. This would
result in greater calcite dissolution in the more
productive areas, and is what we would expect with
higher productivity and diatom production along
the Peru margin and in the equatorial upwelling
systems. These results indicate that rain ratios
change across the EEP, and that these changes
impact the deep ocean, being registered in deep-sea
sediments. So, ballasting of the organic carbon flux
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Fig. 7. The percent calcite preserved difference between the MFI index and the model [Muds] plotted against the organic carbon flux
estimated from the Berger et al. (1988, 1989) (A) and Franc- ois et al. (2002) (B) algorithms.
does not fix the rain ratio to some constant value in
the EEP, and variation in the ratio is both possible
and can be transmitted to depth.
3.7. The rain ratio in the EEP
We can examine the rain ratio variations that
would be compatible with the calcite accumulation
rates and preservation states that we observe in the
EEP by dividing the seabed organic carbon flux we
calculate from the predictive algorithms of Berger
et al. (1988, 1989) and Franc- ois et al. (2002) by the
calcite flux we derive from the accumulation rates
and the MFI index. Fig. 9 illustrates maps of the
rain ratio for the EEP reconstructed with Berger
et al. (1988, 1989) and Franc- ois et al. (2002)
algorithms for organic carbon flux. The regional
distribution of the rain ratio is similar in both maps
despite the differences in underlying assumptions in
each algorithm. The values for the ratio that we
obtain in the open-ocean region are like those found
for comparable water depths in open-ocean settings,
and fit with expectations from sediment traps, and
from modeling water-column and seabed geochemistry (Archer, 1996a, b; Antia et al., 2001). Both rain
ratio maps in Fig. 9 show lowest ratios to the
southwest in the subtropical gyre margin, and
highest values along the equator and closer to the
margin of the Americas. Highest rain ratios (derived
from the Berger et al. (1988, 1989) algorithm)
(Fig. 9(A)) are found in the map along the margin of
Peru. Data for that area are not available with the
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Fig. 8. Map of the difference in percent calcite preserved for the
MFI index and model [MFI model], showing increasing
difference towards the upwelling systems on the equator and
continental margin.
Fig. 9. Rain ratio maps for the Berger et al. (1988, 1989) (A) and
Franc- ois et al. (2002) (B) organic carbon flux algorithms. Data
for the Franc- ois et al. (2002) estimates are a subset of those
available for the Berger et al. (1988, 1989) estimates, so values
cannot be shown in the SE quadrant of the map where surfaceocean productivity is highest.
13
Franc- ois et al. (2002) estimator (Fig. 9(B)). A
quantitative comparison of the rain ratios between
these two organic carbon flux algorithms yields a
consistent relationship even though the scaling for
the rain ratio differs (R2 ¼ 0.88) (Fig. 10). The
Franc- ois et al. (2002)-based rain ratios range in
from about 0.3 to nearly 4, with a trend yielding
values that are roughly 35ths those of the Berger et al.
(1988, 1989)-based ratios. The difference in scaling
can be attributed to the important role that calcite
flux has in ballasting in the Franc- ois et al. (2002)
algorithm, which will reduce the range allowed to
the rain ratio. However, this ballasting component
does not entirely offset rain ratio variation at depth,
and as it is recorded in EEP sediments.
There are many components to our rain ratio
estimates, and each brings its error into our
estimates. The PP values are the result of modeled
satellite chlorophyll concentrations with a reported
uncertainty of o10% (Behrenfeld and Falkowski,
1997). The Behrenfeld and Falkowski (1997) estimates for PP are similar to those from historical
data summarized in Berger et al. (1989). A 10%
change in PP would change our organic carbon flux
and rain ratio estimates also by about 10%
regardless which algorithm we use, Berger et al.
(1988, 1989) or Francois et al. (2002).
Because our sources of error are numerous (e.g.,
errors associated with trapping efficiencies in each
organic carbon attenuation algorithm in Table 2,
Fig. 10. Comparison of the rain ratios from the Berger et al.
(1988, 1989) and Franc- ois et al. (2002) algorithms.
Please cite this article as: Mekik, F., et al., Rain ratio variation in the Tropical Ocean: Tests with surface sediments in the eastern
equatorial Pacific, Deep-Sea Research II (2007), doi:10.1016/j.dsr2.2007.01.010
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estimation of f-ratios, measurement errors associated with each of our proxies, etc.), we cannot
absolutely determine the value for the rain ratio for
any given sample. However, we are able to illustrate
trends in the regional patterns of the rain ratio and
that the rain ratio is strongly variable in the EEP.
Despite using organic carbon flux algorithms
both with and without ballasting in reconstructing
the rain ratio at the seabed, we see a strong gradient
to the rain ratio across the EEP. The scaling of our
rain ratio values depends on the algorithms used for
organic carbon flux estimates, but the range of
values is broad.
Fig. 9 shows a definite link between increasing
surface-ocean productivity and higher rain ratios,
but it also shows two ‘tongues’ of higher rain ratios
that extend into lower-productivity regions. These
are seen in rain ratio values estimated with both
Berger et al. (1988, 1989) and Franc- ois et al. (2002)
algorithms (Fig. 9(A) and (B)). One tongue follows
the South Equatorial Current west of Peru and the
other extends into the Panama Basin. These rain
ratio features lie beneath the Southern Trade Winds
(Li and Philander, 1996) and the isthmus of Panama
easterlies wind stream (QuikScat wind image;
McClain et al., 2002) Panama provides one of the
pathways for the northern Trades to cross from the
Caribbean into the Pacific. These are strong surface
winds that develop especially in the northern
summer months. It is interesting that the region
along the margin of South America, which lies
between the Gulf of Panama and the coast of Peru,
has lower rain ratios in both panels of Fig. 9. This
area is dominated by weaker southerly to onshore
winds in the northern summer months (McClain
et al., 2002; Li and Philander, 1996). The linkage of
rain ratio to easterly winds coming off landmasses
suggests that aeolian supply of micronutrients like
iron might play a role in stimulating diatom
production and driving the rain ratio up. This
mechanism appears to work even where reduced
supply of major nutrients leads to lower overall PP.
Obviously, the data on which the rain ratio
conjecture above is based are limited. However, the
observation may be important, as it would indicate
that significant rain ratio change can happen even in
oceanic regions with moderate PP, and that this
change can be transmitted to the deep ocean. More
research is needed to test these ideas. We currently
lack sufficiently comprehensive sampling to test for
a pattern in detrital sediment accumulation rates
that matches the rain ratio ‘tongues’ or the surface
wind streams. However, we can use the strong
statistical association between detrital component
and organic carbon accumulation rates in the EEP
to infer what the detrital accumulation rate pattern
would look like for the EEP. The organic carbon
data in Fig. 2(D) have features similar to our rain
ratio maps (Fig. 9), with higher values extending
outwards from the continental margin under the
easterly wind paths. Thus, it may be that detrital
material in the sediments reflects the wind paths and
the delivery of key trace elements from the
continents to the surface waters of the EEP.
4. Conclusions
We present the first regional mapping of surface
sediment component accumulation rates in the EEP
based on the 230Th-normalization technique. This
does not depend on averaging sedimentation rates
over a longer sediment interval, and thus time
averaging over thousands of years. So, it is more
likely to represent recent processes in the tropical
Pacific.
We use satellite-derived PP estimates and algorithms for translating these into water-column flux
values to calculate labile organic carbon arriving at
the seabed in the EEP. This, coupled with seabed
calcite flux calculated from accumulation rates and
correction for dissolution, allows us to estimate rain
ratios at depth across the EEP. Also, the sediment
record of calcite preservation allows us to test for a
variable rain ratio signature in the EEP.
We find that calcite preservation patterns across
the EEP are best accounted by variable deep-ocean
rain ratios and highest values for the ratio beneath
higher productivity regions along the equator and
along the margin of the Americas. The rain ratios
that we calculate from algorithm-derived organic
carbon fluxes and reconstructed calcite fluxes also
show this pattern. We find variable deep-ocean rain
ratios in methods with, and without, the inclusion of
ballasting. Regional patterns in the rain ratio are the
same whatever approach we use for calculating
these. Our data indicate that the rain ratio
propagates from the upper ocean into the deep
sea, and leaves an imprint in the sediment record.
The rain ratio variation across the Pacific is likely to
be substantial and must have an impact on carbon
cycling in both the upper and deeper ocean.
We observe higher rain ratios in oceanic areas
under the influence of strong easterly surface winds
that blow from the continent out over the sea. We
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see this in both high- and moderate-productivity
regions. This may indicate the aeolian input of trace
nutrients that enhance diatom production, which is
otherwise Fe- and Si- limited in the EEP.
Lastly, we observe that sedimentary organic
carbon accumulation likely carries a strong allochthonous component because we find no association between organic carbon accumulation rates
and calculated organic carbon fluxes to the seabed,
but we do find a strong linear relationship to detrital
sediment accumulation rates in the EEP. This
indicates that organic carbon accumulation rates
in the EEP do not necessarily provide a good tracer
for past changes in flux of labile organic carbon to
the seabed.
Acknowledgments
Special thanks to Roger Franc- ois (now at the
University of British Columbia, Vancouver, BC) for
allowing us to use his lab at Woods Hole Oceanographic Institution for 230Th analyses and for many
fruitful discussions about ballasting and the rain
ratio. Thanks also to Susan Brown-Leger and Alan
Fleer for assistance with 230Th work at Woods Hole
Oceanographic Institution. This manuscript benefited from thorough and constructive reviews by
Zanna Chase and Guy Munhoven. We are also
grateful to Mark Howland for his help in drafting
some of our figures. Many thanks to core curators
at Oregon State University (June Padman and
Bobbi Conard), Lamont Doherty Earth Observatory (Rusty Lotti-Bond), Rosenstiel School of
Marine Science (Larry Peterson), Scripps Institution
of Oceanography (Warren Smith), University of
Rhode Island (Steve Carey) and the Ocean Drilling
Program for providing us with samples. This work
was supported by NSF Grant no. OCE0326686 and
by a seed Grant from the NASA Michigan Space
Grant Consortium, 2001.
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Please cite this article as: Mekik, F., et al., Rain ratio variation in the Tropical Ocean: Tests with surface sediments in the eastern
equatorial Pacific, Deep-Sea Research II (2007), doi:10.1016/j.dsr2.2007.01.010