Chemical Geology 238 (2007) 180 – 196
www.elsevier.com/locate/chemgeo
Sedimentary redox conditions in continental margin sediments
(N.E. Pacific) — Influence on the accumulation of
redox-sensitive trace metals
J.L. McKay a,⁎, T.F. Pedersen b , A. Mucci c
b
a
Centre GEOTOP, Université du Québec à Montréal, Montréal, Québec, Canada H3C 3P8
School of Earth and Ocean Sciences, University of Victoria, Victoria, British Columbia, Canada V8W 3P6
c
Department of Earth and Planetary Sciences, McGill University, Montreal, Quebec, Canada H3A 2A7
Received 14 February 2006; received in revised form 27 October 2006; accepted 25 November 2006
Editor: D. Rickard
Abstract
Redox conditions in near-surface sediments from the continental margin off western Canada were characterized to determine how
they influence the accumulation of redox-sensitive trace metals. Despite the development of suboxic conditions within millimeters of
the sediment–water interface there is little enrichment of Re, U, Cd or Mo in the upper 10 to 15 cm of the sediment column. This
observation is consistent with the low extractible solid Mn concentrations but high I/Corg ratios that imply that near-surface sediment
redox conditions are poised between Mn and I reduction. In contrast, below ∼15 cm Re, and to a lesser extent Cd and U, are enriched
(e.g., up to 65 ng/g Re) but only in those cores collected from within and below the oxygen minimum zone. Molybdenum concentrations remain low (b1.5 μg/g) in all cores indicating that these near-surface sediments never became fully anoxic.
A combination of low sedimentation rates on the slope (b1 to 6 cm/kyr) and/or intense bioturbation at all locations, including
on the shelf where the sedimentation rate is substantially higher (40 cm/kyr), has maintained suboxic conditions in these sediments.
Suboxic conditions exist despite a high organic carbon flux to the sediment and low bottom water oxygen concentrations, both of
which should favour the development of anoxic conditions close to the sediment–water interface. The enrichment of Re, U and Cd
that is observed occurs well below the sediment–water interface, overprinting and thus potentially compromising any paleo-signal.
Accordingly, caution is necessary when using redox-sensitive trace metal concentrations as paleo-proxies in environments
characterized by low sedimentation rates and/or deep bioturbation.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Redox-sensitive trace metals; Marine sediments; Molybdenum; Rhenium; Paleoceanography
1. Introduction
The accumulation of certain trace metals in sediments
is directly or indirectly controlled by redox conditions
⁎ Corresponding author. Tel.: +1 514 987 4080, fax: +1 514 987 3635.
E-mail address: [email protected] (J.L. McKay).
0009-2541/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.chemgeo.2006.11.008
through either a change in redox state (e.g., Re and U)
and/or speciation (e.g., Mo). Other metals (e.g., Cd) have
a single stable valence in aqueous solution but can be
precipitated in the presence of dissolved sulfide. Observations that certain metals are enriched under specific
redox conditions, such as Mo in anoxic/sulfidic sediments, has led to their use as paleo-environmental proxies
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
(e.g., Rosenthal et al., 1995a; Crusius et al., 1999; Dean
et al., 1999; Yarincik et al., 2000; Adelson et al., 2001;
Pailler et al., 2002; Ivanochko and Pedersen, 2004).
Changes in the intensity of the oxygen minimum zone on
the northeastern margin of the Pacific, for example, have
been inferred from variations in the sedimentary concentrations of redox-sensitive trace metals (Dean et al.,
1997; Zheng et al., 2000; Nameroff et al., 2004; McKay
et al., 2005).
The application of redox-sensitive trace metals as
paleo-proxies requires an understanding of their geochemical behaviours in a wide variety of modern environments. Anoxic environments such as the Black Sea
and Cariaco Basin have been fairly well studied (e.g.,
Calvert, 1990; Dean et al., 1999; Yarincik et al., 2000).
In comparison, detailed studies of the more typical continental margin settings are relatively limited (Crusius
et al., 1996; Dean et al., 1997; Morford and Emerson,
1999; Nameroff et al., 2002; Sundby et al., 2004). The
goal of this study is to provide further insight on the sedimentary geochemistry of redox-sensitive trace metals
(Re, U, Cd and Mo) in a modern continental margin
setting characterized by low to moderate sedimentation
rates (b 1 to 40 cm/kyr).
The study area is located at the northern end of the
California Current System off Vancouver Island, British
Columbia, Canada (Fig. 1). This region sits within a
transitional zone where the eastward flowing Subarctic
and North Pacific currents split into the northward
flowing Alaska Current and southward flowing Cali-
Fig. 1. Study area off the west coast of Vancouver Island, British
Columbia, Canada (inset) showing the locations where multicores and
box cores were collected. Exact station locations and water depths are
provided in Table 1.
181
fornia Current. The area immediately west of Vancouver
Island is typified by high primary productivity (up to
400 gC/m2/yr; Antoine et al., 1996) related to seasonal
(late spring to early fall) wind-induced upwelling of
nutrients (Huyer, 1983). The region is also characterized
by a relatively intense oxygen minimum zone (OMZ)
that is most pronounced between 700 and 1300 m (0.3 to
0.5 ml/l O2). During periods of upwelling this oxygenpoor water is transported onto the shelf (Mackas et al.,
1987). Low bottom water oxygen concentrations in combination with a high organic carbon flux to the sediment
should normally lead to the development of anoxic
conditions within centimeters of the sediment–water
interface and result in the accumulation of certain redoxsensitive trace metals (e.g., Re, U, Cd and Mo). However, the surface and near-surface sediments on the Vancouver Island continental margin are not enriched in
these trace metals. In this paper we investigate the early
redox conditions that characterize these sediments and
propose why redox-sensitive trace metals are not readily
sequestered during early diagenesis.
2. Methods
Box cores (bc, 0.12 m2 Ocean Instruments Mark II
box corer) and multicores (mc, Bowers and Connelly
multicorer with eight 10 cm-diameter core barrels) were
collected from six stations on the continental margin off
the west coast of Vancouver Island along a transect that
extended from the shelf down the slope to a water depth
of 1750 m (Fig. 1). Core 01 was collected from a
depression on the shelf (120 m). Core 04 was collected
from the upper slope above the OMZ (407 m). Core 06
is from the upper edge of the OMZ (720 m), Core 09
from within the OMZ (920 m), and Core 02 from just
below the OMZ (1340 m). Core 05 was collected from
the lower slope and is located well below the OMZ
(1750 m). Table 1 provides the location, water depth and
bottom water oxygen concentration for each station as
well as a description of each core.
A sedimentation rate for each core was established by
assigning an age of 0 years to the core top and obtaining
an age estimate from near the core base. Linear sedimentation was assumed between the two assigned ages.
When possible, a mixed assemblage of planktonic foraminifera was radiocarbon dated by accelerator mass
spectrometry (AMS). However, this was not always feasible given the scarcity of carbonate microfossils, in
which case bulk organic carbon was dated (e.g., cores 01
and 02). Dating bulk organic carbon typically yields
older than expected ages due to the presence of old,
recycled organic matter. To validate the use of bulk
182
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Table 1
General data for sampling locations and core descriptions
[O2] in OLW a (ml/l)
OPD c (mm)
Core description
120
407
2.4
1.0
4 to 5
–
126° 52.68' W
720
0.4
2 to 3
126° 53.44′ W
127° 18.57′ W
127° 33.12′ W
920
1340
1750
0.3
0.4 b
1.2
5
6 to 7
–
38 cm olive green mud
∼ 6 to 8 cm olive green muddy sand underlain
by a gray clay
∼ 18 cm olive green sandy mud underlain
by a gray clay
40 cm olive green mud
20 cm olive green mud
48 cm olive green mud
Core
Latitude
Longitude
01
04
48° 45.95′ N
49° 00.71′ N
125° 29.57′ W
126° 49.82′ W
06
48° 58.73' N
09
02
05
48° 54.76′ N
49° 12.81′ N
49° 07.91′ N
a
b
c
Water depth (m)
Oxygen concentration in overlying water (OLW) obtained from bottle casts ∼15 m above bottom.
Overlying water oxygen concentration for Station 02 was measured at 1240 m, not 1340 m.
Oxygen penetration depth determined by voltammetric micro-electrode profiling.
organic carbon ages, 14C measurements of coexisting
foraminifera and organic carbon were obtained for one
sample taken from Core 09. For cores 06 and 04 radiocarbon measurements are not available and sedimentation rates were estimated by assuming that the contact
between the green mud and greenish-gray clay, which is
observed in both cores, is the same age as the mud-clay
contact observed in Piston Core JT96-09 from Site 09
(i.e., ∼ 11.2 kyr; McKay, 2003). These ages are minimum ages because we cannot rule-out the possibility of
erosional events and/or hiatuses during the deposition of
the Holocene sediments.
The extent of bioturbation was estimated using excess
210
Pb data. The analytical method of Eakins and Morrison
(1978) was employed, which determines the 210Pb concentration by measuring the content of its 210Po granddaughter via alpha counting.
Oxygen concentrations in the overlying bottom water
were determined by Winkler titration. High-resolution
profiles of porewater O2 and HS− were measured simultaneously on multicores recovered from stations 01, 06,
09 and 02 using a solid-state voltammetric microelectrode similar to that described in Brendel and Luther
(1995). Each core was mounted upright and immersed
in an ice-water slurry to maintain its temperature during
profiling. Electrode calibration procedures and working
conditions are described in detail in Luther et al. (1998).
Upon recovery, the box cores were placed upright in
a large glove box purged by a continuous flow of nitrogen to minimize sediment oxidation (Edenborn et al.,
1986). The sediments were sub-sampled typically every
0.5 cm over the first centimeter, every centimeter over
the next 5 to 10 cm, and at 2 to 5 cm intervals in the
lower portion of the core. As each sampling interval was
sequentially exposed, sediments were transferred to
porewater squeezers. Mini-cores (i.e., a 10 cc syringe
plunger inserted into a 13 ml polyethylene screw cap test
tube that has the distal end cut off) were also taken at the
rate of two per depth interval and frozen immediately in
order to maintain the redox conditions of the sediment.
These samples were used for acid volatile sulfide (AVS)
determinations. Separate cores (box cores or multicores
depending on availability) were sampled at a 1 to 2 cm
resolution for major, minor and trace element analysis of
the sediment. The concentration of pyrite and various
biogenic components (i.e., organic carbon, carbonate
and opal) was also measured. Prior to analysis these
samples were freeze-dried and hand ground, except for
the sandy sediments of Core 04 which were pulverized
in a Tema WC disc mill.
Porewaters were extracted from cores 01, 04, 06 and
09 using Reeburgh-type squeezers (Reeburgh, 1967),
modified to filter the water through a 0.45 μm Type HA
Millipore filter as it passed directly into a 50 cc syringe.
Dissolved ammonium concentrations were determined
on-board ship within a few hours of their extraction. The
remaining porewater was divided into two aliquots. One
fraction was stored untreated in pre-washed polyethylene bottles for sulfate analyses and the second was
acidified with a 1% equivalent volume of Seastar concentrated HCl for nitrate and total dissolved Fe and Mn
determinations.
Ammonium concentrations were measured by the
conductivity method of Hall and Aller (1992). The
analytical precision of this method is better than 5%.
Porewater sulfate concentrations were determined by
ion chromatography (Dionex, 1986) after a 100-fold
dilution in distilled water. The method was calibrated
using IAPSO Standard Seawater and has a reproducibility of ±1%. Nitrate levels were measured spectrophotometrically according to the method described by
Strickland and Parsons (1972) with a reproducibility of
better than 3%. Total dissolved Fe and Mn were measured by flame atomic absorption spectrophotometry
(AAS) on a Perkin–Elmer Model 3100 AAS. Reproducibility of these measurements is better than ± 5%.
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
The concentration of acid volatile sulfide (AVS),
which includes mackinawite, poorly crystallized greigite and dissolved sulfide species, was determined
according to the method of Hsieh and Yang (1989) as
modified by Gagnon et al. (1995). This involved reacting the wet sediment for 3 h with a mixture of
9 N HCl and 20% SnCl2 at room temperature in a N2flushed and sealed vial. The evolved H2S was then
trapped by diffusion into a zinc chloride solution and
titrated iodometrically. Amorphous Fe and Mn oxides
were extracted from the sediment using a buffered
ascorbic acid solution (Kostka and Luther, 1994). Total
Fe and Mn oxides (i.e., amorphous + crystalline oxides)
were selectively dissolved by a citrate–dithionate–
bicarbonate (CDB) buffered extraction (Lucotte and
d'Anglejan, 1985). Pyrite content was measured according to the method described by Lord (1982). A
known weight of freeze-dried sediment was sequentially treated with a citrate–bicarbonate–dithionate
solution and hydrofluoric–boric acid to remove all Fe
compounds except pyrite. The remaining Fe, assumed
to be only in the form of pyrite, was then dissolved in
concentrated nitric acid and the Fe concentration of the
final solution was determined by atomic absorption
spectrophotometry (AAS) with a reproducibility of 5%.
Total carbon (Ctot) and total sulfur (Stot) were measured using a Carlo Erba NA-1500 elemental analyzer.
Precision and accuracy were determined on the basis of
repeated analyses of two National Research Council of
Canada (NRC) standards (PACS-1 and MESS-1). The
measurements have relative standard deviations (RSD,
1σ) of 3% and 6% for C and S respectively, and the
accuracy is within 7% of the certified values for both
elements. Carbonate carbon (Ccarb) was determined by
coulometry and repeat analyses (n = 141) of a CaCO3
standard yielded a mean value of 11.9% with a RSD of
1%. The percent organic carbon (Corg) was calculated as
the difference between total carbon and carbonate
carbon contents, and has an aggregate error (as RSD)
of ∼ 4%. Biogenic silica was measured using the
Na2CO3 dissolution method of Mortlock and Froelich
(1989). The RSD determined for two in-house standards
SNB and JV5, which contain approximately 28% and
11% opal respectively, is 4%. However, the method is
less precise when opal concentrations are b 10%, most
probably due to the dissolution of volcanic glass and
clays. The Al and I contents were determined by X-ray
fluorescence spectrometry and have a RSD of ∼5%.
The concentrations of Re, U, Cd and Mo were measured by isotope-dilution inductively-coupled plasma
mass spectrometry. Sample preparation involved adding
known amounts of isotopically-enriched spike solutions
183
to 15–20 mg of powdered sediment. Samples were then
microwave digested in a mixture of concentrated, ultrapure HNO3, HCl and HF (0.8 ml, 0.1 ml, and 0.2 ml,
respectively). The digests were evaporated on a hotplate
overnight and then re-dissolved in 5 N HCl. Precision of
these analyses was evaluated by analyzing two standards, the University of British Columbia Saanich Inlet
sediment standard (SNB) and the NRC standard MESS2, with each batch of samples. SNB is characterized by
high concentrations of redox-sensitive trace metals (i.e.,
average values of 5.61 ng/g Re, 4.31 μg/g U, 4.86 μg/g
Cd, and 54.50 μg/g Mo; n = 35), consistent with the
anoxic depositional environment. In comparison,
MESS-2 is typified by lower trace metal concentrations
(i.e., average values of 2.88 ng/g Re, 2.68 μg/g U,
0.24 μg/g Cd, and 2.37 μg/g Mo; n = 30). The RSDs
(1σ) determined using SNB are 9% for Re, 12% for U,
and 6% for both Cd and Mo. The RSDs determined
using MESS-2 are slightly higher except for that of U
(13%, 9%, 13% and 9%; Re, U, Cd and Mo respectively). The accuracy of Cd and Mo measurements,
assessed using MESS-2 (certified values of 0.24 μg/g
Cd and 2.85 μg/g, Mo), are b 1% (Cd) and 17% (Mo).
There are no certified MESS-2 values for Re and U.
3. Results
3.1. The cores
Core descriptions are provided in Table 1. The homogeneous olive green muds, sandy muds and muddy
sands observed in these near-surface sediment cores are
typical of Holocene deposits off Vancouver Island
(Bornhold and Barrie, 1991). The greenish-gray clay
found at the base of cores 04 and 06 is most probably
Pleistocene in age and of glaciomarine origin (Bornhold
and Yorath, 1984; Bornhold and Barrie, 1991). This
conclusion is consistent with radiocarbon dating of
Piston Core JT96-09, collected from Site 09, which
yields a date of ∼ 11.2 calendar kyr for the contact between the green mud and gray clay (McKay, 2003).
Radiocarbon data and estimated sedimentation rates
are given in Table 2. The radiocarbon date obtained
for bulk organic carbon in Core 09 (9920 ± 40 14C years)
is similar to the age determined by dating planktonic
foraminifera found in the same depth interval (9760 ±
110 14C years). This result suggests that bulk organic
matter can be used for dating these near-surface sediments. However, the use of bulk organic carbon dates is
only valid for Holocene sediments that contain little
terrestrial organic matter, and not for Late Pleistocene
sediments that contain up to 70% old terrestrial organic
184
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Table 2
Radiocarbon data and estimated sedimentation rates
Core
Water depth (m)
Material dated
Sample depth (cm)
14
Calendar age (yrs) b
Sedimentation rate (cm/kyr)
01
04
06
09
120
407
720
920
02
05
1340
1750
Bulk organic matter
–
–
Planktonic foraminifera a
Bulk organic matter
Bulk organic matter
Planktonic foraminifera
36.5
–
–
47.5
39.0
19.0
47.0
1780 ± 30
–
–
9760 ± 70
9920 ± 40
3605 ± 35
9960 ± 50
923
–
–
10026
10281
3003
10290
39.5 c
0.7 d
1.6 d
4.7 c
3.8 c
6.3 c, 2.1 d
4.6 c
C Age (yrs)
a
Planktonic forams were obtained from Piston Core JT96-09 that was collected from the same location.
AMS 14C ages were converted to calendar years using Calib 4.3 (Stuiver et al., 1998) and assuming a reservoir age of 800 years.
c
Sedimentation rates were determined using AMS 14C ages.
d
Sedimentation rates were estimated assuming that the contact between the Holocene muds and Pleistocene gray clay is at 11.2 calendar kyr. This
contact occurs at 8 cm in Multicore 04, 18 cm in Multicore 06 and 24 cm in Multicore 02. Note, this contact is not observed in Multicore 02 so its
depth has been determined using the piston core.
b
matter (McKay et al., 2004). It should also be noted that
sedimentation rates determined from radiocarbon data
could be overestimated due to the bioturbation of younger
material downward. The generally low Holocene sedimentation rates (i.e., b 10 cm/kyr except at Station 01)
result from trapping of river-borne detritus within the
fjords that are ubiquitous along the coastline (Bornhold
and Yorath, 1984). In addition, strong waves and currents
affect the shelf and upper slope (Bornhold and Yorath,
1984), thus explaining the very low sedimentation rates at
Stations 04 and 06. In contrast, Station 01, which sits
within a depression on the shelf, is not affected to the same
degree by the winnowing action of waves and currents and
thus has a much higher sedimentation rate (40 cm/kyr).
Total 210 Pb concentrations in surface sediments
range from 24.5 dpm/g in Core 01 up to 58.8 dpm/g in
Core 02. Supported 210 Pb produced in situ by the decay
of 226 Ra ranges from 1.8 to 2.8 dpm/g, similar to that in
sediments from the Washington State continental margin just south of Vancouver Island (Carpenter et al.,
1981). Unsupported or excess 210 Pb (210 Pbxs), which is
scavenged from the water column, is found as deep as
28 cm in Core 01, 19 cm in Core 09 and 11 cm in Core
02. This is much deeper than expected based on the
estimated sedimentation rates (b 1 to 40 cm/kyr) and the
short half-life of 210 Pb (22.3 years). A plot of ln 210 Pbxs
versus depth clearly indicates that the upper 15 cm of
Core 01 has been recently mixed (i.e., little change in
ln 210 Pb xs with depth; Fig. 2), most probably by
bioturbation. In contrast, Core 02 exhibits a linear
decrease of ln210 Pbxs with depth (Fig. 2), as expected
from the exponential radioactive decay law. The
sedimentation rate calculated using these 210 Pb data
is, however, unrealistically high (82 cm/kyr) considering that the Holocene mud is only 25 cm thick in Core 02,
which yields a sedimentation rate of ∼2 cm/kyr, a value
more consistent with the one derived from 14C dating
(∼6 cm/kyr). The simplest explanation is that 210Pb has
been mixed into the sediment via bioturbation, although
not as deeply or completely as in Core 01. The argument is
similar for Core 09, the ln210Pbxs profile results from a
combination of the exponential decay of 210Pb and a
similarly exponential decrease of bioturbation intensity
with depth (Middelburg et al., 1991). By mixing younger
sediment downwards, bioturbation will also decrease the
14
C age of sub-surface deposits, which as noted above, can
lead to overestimated sedimentation rates.
Fig. 2. Downcore profile of ln 210Pbxs in cores 01mc, 02mc and 09mc.
The occurrence of 210Pbxs at depths greater than a few centimeters
suggests that all cores, particularly Core 01, have been recently and
deeply bioturbated.
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
185
Fig. 3. Porewater profiles (sulfate, nitrate, ammonium, Mn2+ and Fe2+) for Core 01bc (a and b), Core 04bc (c and d), Core 06bc (e and f ), and Core
09bc (g and h). The dashed lines indicate the position of the contact between the Holocene green mud and Pleistocene gray clay in cores 04 and 06.
Note, that the scales of the x-axes for the graphs showing Mn2+ and Fe2+ concentrations are not all the same.
186
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Fig. 4. Concentrations of ascorbate-extractible (ASC) Fe and Mn, citrate–dithionate–bicarbonate extractible (CDB) Fe, acid volatile sulfide (AVS)
and pyrite in Core 01bc (a and b), Core 04bc (c and d), Core 06bc (e and f ) and Core 09bc (g and h). The dashed lines indicate the position of the
contact between the Holocene green mud and Pleistocene gray clay in cores 04 and 06.
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
3.2. Porewaters
Micro-electrode analyses of cores 01, 06, 09 and 02
(B. Sundby, unpublished data) revealed that the oxygen
penetration depth is less than 1 cm (Table 1). Below this
oxic zone, traces (i.e., b 10 μM) of HS− appear over the
next 4 to 5 cm.
Porewater concentration profiles for sulphate, nitrate
and ammonium, as well as Mn(II) and Fe(II) in cores 01,
04, 06 and 09 are shown in Fig. 3. With the exception of
187
Core 01, there is no discernible depletion of porewater
sulfate with depth. In contrast, nitrate concentrations
decrease rapidly in the upper 1 cm of all cores. The
strongest dissolved ammonium gradient, a surrogate
tracer of suboxic and anoxic diagenesis, is observed in
Core 01. In comparison, only small increases in dissolved ammonium concentrations with increasing depth
are observed in the other cores. Core 01 also has the
highest dissolved Mn(II) and Fe(II) concentrations and
both metals exhibit a sub-surface maximum at ∼ 5 cm
Fig. 5. a) The downcore concentrations of total sulfur, b) organic carbon, c) carbonate and d) opal.
188
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
(Fig. 3b) that most probably corresponds to the reductive dissolution of reactive Mn and Fe oxides (i.e.,
the oxic–suboxic boundary). In the other cores, dissolved manganese and iron concentrations are generally
low (b20 μM).
3.3. Solid sediments
The ascorbate (ASC) extraction is believed to be
selective for amorphous Fe oxides whereas the citrate–
dithionate–bicarbonate buffered (CDB) extraction will
dissolve both the amorphous and crystalline Fe oxides
(Kostka and Luther, 1994). Accordingly, the ascorbateextractible Fe (FeASC) concentrations are smaller than
the CDB-extractible Fe (FeCDB) and better define the
authigenic metal enrichment at, or near, the oxic–suboxic boundary. In general, FeCDB concentrations show
only minor variations with depth in the Holocene mud,
but increase sharply in the Pleistocene gray clay found at
the bottom of cores 04 and 06 (Figs. 4c and e). The increase of FeCDB in the gray clay is correlated to the
presence of pyrite and most probably reflects the
extraction of Fe from this sulfide or phases that formed
from it during freeze-drying. In all cores, the MnASC and
FeASC profiles exhibit a slight enrichment near the sediment–water interface and relatively constant values at
depth (Figs. 4a, c, e and g). The low extractible-Mn
values (b0.5 μmol/g) may reflect a lack of supply of
detrital Mn oxides and/or the loss of Mn(II) to the
overlying water column. Acid volatile sulfide (AVS) and
pyrite concentration profiles differ substantially from
core to core. The Holocene sediments in Core 01 have
the highest AVS and pyrite concentrations (∼ 15 and
70 μmol/g, respectively; Fig. 4b). In cores 04 and 06,
AVS and pyrite concentrations are low in the near-surface Holocene sediments and increase abruptly within
the Pleistocene gray clay (Figs. 4d and f). In Core 09,
AVS and pyrite contents are low throughout most of the
core, with pyrite increasing only slightly near the base of
the core (Fig. 4h).
Total sulfur (Stot) concentration in surface sediments
range from b0.1 wt.% in Core 04 to 0.4 wt.% in cores 02
and 09. The gray clay, which occurs in lower portions of
cores 04 and 06, is characterized by relatively higher Stot
concentrations than the Holocene sediments in these
cores (Fig. 5a), consistent with the relatively high AVS
and pyrite concentrations. There is also a substantial
increase in Stot below ∼ 20 cm in Core 01 that is unrelated to a change is sediment lithology, but corresponds
well with the increase in pyrite content.
Organic carbon (Corg) contents are highest in the
surface sediments of cores 02 and 09 (3.4 and 3.1 wt.%,
respectively) and decrease with depth (Fig. 5b). Lower
concentrations (1.6 to 2.1 wt.%) are observed in cores
01, 05 and 06 and, with the exception of 06, decrease
only slightly with depth (Fig. 5b). The abrupt decrease
in Corg observed at ∼18 cm depth in Core 06 occurs at
the contact between the Holocene mud and Pleistocene
gray clay. The lowest Corg content (0.5 to 0.6 wt.%) is
found in the sandy sediments of Core 04. Based on its
δ13C signature, the organic matter found within the
Holocene sediments is predominantly marine (b 22%
terrestrial material) while the Pleistocene gray clay
contains significantly more terrestrial organic matter
(N 39%; McKay, 2003). The organic carbon mass
accumulation rate, calculated as the product of the
linear sedimentation rate, the dry bulk density and the
organic carbon concentration, is lowest in cores 02, 05
and 09 (≤ 0.1 g/cm2/kyr) and up to an order of magnitude higher in Core 01 (0.3 to 0.6 g/cm2/kyr). The
CaCO3 content is low in all cores, ranging from b1.2 wt.%
in the surface sediments up to 3% in the lower portions
of cores 05, 06 and 09 (Fig. 5c). Opal concentration
ranges from b 3 up to 11 wt.% in surface sediments (cores
04 and 01, respectively) and exhibits little change
downcore (Fig. 5d).
Iodine concentrations in surface sediments range from
104 to 968 μg/g. The I/Corg ratios in the surface sediments
Fig. 6. I/Corg ratios in near-surface sediment cores.
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
range from 120 × 10− 4 in Core 01 up to 500 × 10− 4 in Core
05, and decrease gradually with depth in all cores (Fig. 6).
Downcore trace metal profiles are provided in Fig. 7.
Data are presented as concentrations, rather than metal/
Al ratios, because the Al content was not measured for
every sample, and the metal concentration data therefore
yield a finer resolution. This approach is acceptable
because the Al content of the Holocene sediments and
189
the gray clay that underlies the Holocene sediments in
cores 04 and 06 are similar (6.2 to 8.8% Al), and thus the
downcore concentration and metal/Al profiles do not
differ significantly. Furthermore, no attempt was made
to calculate the authigenic trace metal fraction because
the lithogenic concentrations of these metals are poorly
constrained. Rhenium concentrations are relatively low
(2 to 4 ng/g) in the surface sediments of all cores. There
Fig. 7. Downcore concentration profiles of sedimentary Cd, Re, U and Mo in cores 01mc, 04mc, 06bc, 09mc, 02mc and 05bc (figures a to f,
respectively). The dashed line in Fig. 6b and c indicate the position of the contact between the Holocene green mud and Pleistocene gray clay that
occurs in cores 04 and 06. To facilitate the plotting of the data, Re concentrations have been divided by 10 and Cd concentrations have been multiplied
by 10.
190
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
is no change in Re concentration with depth in cores 01,
04 and 06 (Figs. 7a to c). In contrast, cores 09, 02 and 05
exhibit sub-surface Re concentration peaks (Figs. 7d
to f) with the highest Re concentration being measured
in Core 09 (65 ng/g). Interestingly, the depth at which
Re enrichment is first observed in cores 09 and 02
corresponds to the depth at which the concentration of
excess 210Pb becomes undetectable (i.e., the base of the
recently bioturbated zone). Uranium concentrations are
low (b 2 μg/g) in the upper 10 to 15 cm of all cores
(Fig. 7). Below this depth in Core 09, and to a lesser
extent in cores 02 and 05, U enrichment (up to 5.8 μg/g)
is observed (Figs. 7d–f). Cadmium concentrations are
also low in the surface sediments (b 0.1 to 0.2 μg/g), and
increase slightly at depth in cores 09, 02 and 05 (Figs. 7d
to f). Once again, the greatest enrichment occurs in Core
09 (up to 1.0 μg/g). The Pleistocene gray clay found at
the base of cores 04 and 06 is also characterized by
higher Cd concentrations (Figs. 7b and c). The Mo
content of surface sediments ranges from 0.4 to 0.9 μg/g
(Fig. 7). There is a very slight increase in Mo
concentration with depth in all of the cores, but there
is no evidence of a substantial enrichment above typical
lithogenic concentrations (∼ 0.6 μg/g for terrigenous
sediments derived from Vancouver Island; Morford
et al., 2001) in any of the cores.
4. Discussion
4.1. Sedimentary redox conditions
Bottom water oxygen concentrations were relatively
low at all stations (0.3 to 2.4 ml/l; Table 1). Low oxygen
concentrations, combined with the relatively high
organic carbon content of the sediment, result in a thin
oxic layer below the sediment–water interface (i.e.,
oxygen penetration depths of b 1 cm) within which
denitrification occurs. At Station 01, a bottom water
NO3− concentration of 30 μmol/l was measured,
compared to only 5.7 μmol/l in the porewater of the 0
to 0.5 cm interval below the sediment–water interface.
This intense denitrification should support a strong flux
of NO3− into the sediment. The rate of denitrification,
evaluated from a linear fit between the NO3− concentration in the overlying bottom water and the depth of nearzero NO3− concentration (∼ 0.75 cm), is 0.27 mmol/m− 2
d− 1. No attempt was made to evaluate denitrification
rates at the other stations because the corresponding
bottom water concentrations were not measured.
Nevertheless, the strong porewater concentration gradient between 0.25 and 0.75 cm below the sediment–
water interface in cores 04 and 06 also suggests that
intense denitrification is occurring at these sites. No
results are available for Station 09.
Under oxic conditions iodate (IO3−) is adsorbed by
marine organic matter (Price and Calvert, 1977).
However, under suboxic conditions IO3− is reduced to I−
and this species is not absorbed by marine organic matter
(Price and Calvert, 1977). Thus, sediments that accumulate in oxygenated environments typically have I/Corg
values of N200 × 10− 4 whereas sediments in suboxic
and anoxic settings are characterized by low (typically
b20 × 10− 4) I/Corg values (Price and Calvert, 1977;
Francois, 1987). Surface sediments in all cores from the
Vancouver Island continental margin, with the exception
of Core 01, are characterized by I/Corg ratios that exceed
200 × 10− 4 (Fig. 6) suggesting that surface sediments are
sufficiently oxygenated to allow adsorption of iodine
species onto organic matter. Iodine is then lost relative to
carbon during diagenesis (Price and Calvert, 1973, 1977;
Francois, 1987) resulting in the observed decrease in I/
Corg ratios with depth (Fig. 6). The low I/Corg ratios
(b 124 × 10− 4) in surface sediments of Core 01 cannot be
attributed to the presence of substantially more terrigenous organic matter, which does not adsorb dissolved
iodine species (Malcolm and Price, 1984). Rather, the low
I/Corg ratios reflect more reducing conditions possibly
related to the high organic carbon mass accumulation rate
(∼0.6 g/cm2/kyr) at this site and correspondingly higher
oxygen demand and/or periods of lower bottom water
oxygen associated with upwelling events.
According to the current paradigm, oxic degradation is
followed by manganese and iron oxide reduction within
the suboxic zone (Froelich et al., 1979). The distribution
of extractible solid Mn and Fe phases (MnASC, FeASC and
FeCDB; Fig. 4) attests to the accumulation of authigenic
Mn and Fe oxides in the surface oxic layer of the sediment
and their reductive dissolution a few millimeters to
centimeters below the sediment–water interface. However, with the exception of Station 01, little Mn(II)
accumulates in the porewater and only minor concentration gradients are established that could sustain a strong
flux of Mn(II) towards the surface. In the absence of
sediment trap data, we can only assume that little
dissolved Mn accumulates in these sediments because
there is a limited flux of Mn oxides to the sediment and/or
dissolved Mn(II) is lost to the overlying water column. In
contrast, notable Fe(II) porewater gradients are observed
at all stations. The gradients are sharpest within the first
centimeter at stations 01 and 06 (Figs. 3b and f), and in the
2 to 3 cm interval at Station 09 (Fig. 3h). This further
constrains the thickness of the oxic zone and reveals that
there is rapid oxidation of porewater Fe(II) once it enters
the thin oxic layer. The differential behaviour of Mn and
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Fe may be the consequence of the low dissolved oxygen
levels in the overlying waters and possibly the slower
oxidation kinetics of Mn(II).
Under anoxic conditions, organic matter oxidation
proceeds via sulfate reduction. There are no direct sulfate
reduction rate measurements available for the study area,
and [SO42−] gradients obtained during this study are too
small to allow reliable flux and rate determinations.
Porewater sulfate depletion is only noticeable in Core 01
(Fig. 3a) and a progressive accumulation of authigenic
sulfide minerals (i.e., AVS and pyrite) with depth is
readily observed (Fig. 4b). At the other stations, AVS and
pyrite concentrations in the Holocene sediments are
generally low and invariant with depth. In contrast, the
Pleistocene gray clay observed in cores 04 and 06 is
characterized by relatively high concentrations of sulfide
minerals (Figs. 4d and f) and correspondingly high FeCDB
values (Figs. 4c and e). This gray clay also contains
significant amounts of porewater ammonium which attest
to the prevailing reducing conditions (strongly suboxic or
anoxic), but the abrupt increase in pyrite at the Holocene–
Pleistocene contact and its rather uniform distribution
within the gray clay leads us to believe that most of the
pyrite is inherited (i.e., not the product of present-day
redox conditions). We know that between ∼13.5 and
12.6 kyr, as the gray clay was being deposited, the OMZ
off Vancouver Island was more intense, and as a result
near-surface sediments were anoxic (McKay et al., 2005).
The AVS and pyrite within the gray clay in cores 04 and
06 were most probably formed at this time.
In summary, the near-surface sediments deposited on
the Vancouver Island continental margin are characterized by a thin oxic zone (b1 cm). Below this oxic layer
sediments become suboxic, resulting in the reductive
dissolution of Mn and Fe oxides. With the exception of
Station 01, there appears to be little net sulphate reduction occurring, and even at Station 01 sulphate
reduction and sulfide precipitation is limited. It is clear
that these sediments do not become fully anoxic despite
their high organic carbon content and the low oxygen
concentrations in the overlying water, both of which
should favour the development of anoxic conditions.
This observation directly influences the sediment's ability to scavenge redox-sensitive trace metals such as Re,
U, Cd and Mo.
4.2. Accumulation of redox-sensitive trace metals
The diffusion of certain trace metals (Re, U, Cd and
Mo) into the sediment from the overlying water column,
and their subsequent authigenic accumulation is controlled by: i) in situ redox conditions, and ii) the
191
porewater concentration gradients generated by their
differing geochemical behaviours across redox boundaries. In general, when the oxic zone is thin (i.e., shallow
oxygen penetration depths) and, in turn, the suboxic and
the anoxic redox boundaries are shallow, concentration
gradients are steeper and the flux from the water column
into the sediment is enhanced. These conditions commonly occur where the organic matter flux to the seafloor
is high and/or oxygen concentration in the bottom water
is low. Alternatively, it is also possible to build-up high
metal concentrations when the suboxic and anoxic redox
boundaries are quite deep, as long as the boundaries
remain nearly stationary for a long period of time. This
scenario is encountered when sedimentation rates are
relatively low, but the oxidant demand is high.
Rhenium is a conservative element in seawater and
its concentration is typically low in oxic marine sediments (≤ 0.5 ng/g; Boyko et al., 1986; Koide et al.,
1986) but may be as much as 300 times higher in suboxic and anoxic deposits (Koide et al., 1986; Ravizza
et al., 1991; Colodner et al., 1993; Crusius et al., 1996;
Morford and Emerson, 1999). Enrichments of this magnitude are the result of Re diffusion from the overlying
water column into the sediment and its fixation under
suboxic conditions. It is typically assumed, although not
yet proven, that this involves reduction from Re(VII)
to Re(IV) and precipitation, possibly as ReO2 (Crusius
et al., 1996). The Re concentration in the upper 10 to
15 cm in all of the Vancouver Island continental margin
cores is only slightly elevated (1.6 to 4.2 ng/g) when
compared to the typical concentration in oxic sediments
(≤0.5 ng/g). Although sediments become suboxic within 1 cm of the sediment–water interface they are not
sufficiently reducing to induce significant Re reduction
and precipitation at these shallow sub-surface depths.
This is in contrast to results from the Washington State
continental margin where Re enrichment is observed in
surface sediments when the oxygen penetration depth is
b 1 cm (Morford et al., 2005). According to Thomson
et al. (1993), Re reduction occurs subsequent to I reduction. Rhenium reduction and accumulation would
therefore not be anticipated in the near-surface sediments
off the Vancouver Island since their generally high I/Corg
ratios indicate that iodine is present in its oxidized form.
Nevertheless, below ∼ 15 cm in cores 09, 02 and 05,
there is a substantial Re enrichment (Figs. 7d to f). We
hypothesize that this reflects ongoing Re diffusion into
the sediment followed by its reduction and precipitation
at depth. Unfortunately, no porewater Re concentration
data are available to permit a direct assessment of this
hypothesis. Instead, a simple steady-state linear-diffusion model is applied to determine whether or not the
192
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
enrichments could reflect ongoing authigenesis. We use
Fick's First Law (Eq. (1)) to compute the downward
diffusion flux (J) for Core 09.
J ¼ ðU=FÞDb ðdc=dzÞ
ð5Þ
The bulk in situ diffusivity (Db) of the perrhenate ion is
assumed to be similar to that of other oxyanions (i.e.
∼ 6 × 10− 6 cm2 s− 1; Li and Gregory, 1974). The porosity
(Φ) is assigned a value of 0.8 (i.e., average value for Core
09) and the formation factor (F) is taken to be 1.3
(Manheim, 1970), a value typical of silty clays. The
concentration gradient (dc / dz) is assumed to be linear
and is computed as the quotient of the difference in concentration between the bottom water (7.8 ng/kg; Anabar
et al., 1992) and a precipitation point 20 cm below the
sediment–water interface where the porewater Re concentration is assumed to be 0. This yields an estimated
flux (J ) of 45 ng cm− 2 kyr− 1. The mass accumulation
rate (MAR) of the sediments in Core 09, computed as the
product of the linear sedimentation rate (∼ 5 cm/kyr) and
the dry bulk density [(1 − Φ) times the grain density of
2.5 g cm− 3], is 2.5 g cm2 kyr− 1. The ratio of J/MAR
represents the steady-state concentration of Re that could
be added to the sediments via downward diffusion to a
fixation depth of 20 cm, and in this example is equal to
∼ 18 ng/g. This amount is of the same order of magnitude
as the observed Re concentration at 20 cm depth in Core
09 (i.e., 22 ng/g). A slightly lower value (∼ 14 ng/g) is
obtained if we assume that the porewater Re concentration does not fall to zero, but to some background level
(e.g., ∼ 1.86 ng/kg = 10 pM; the value measured for similar suboxic sediments from the continental margin off
Washington State; Morford et al., 2005). This admittedly
crude and assumption-dependent result suggests that
downward diffusion and sub-surface precipitation could
account for the Re enrichment in these deposits. On the
other hand, the degree of enrichment is very sensitive to
the sedimentation rate. Core 01, which has a much higher
sedimentation rate compared to Core 09 (39.5 versus
4.7 cm/kyr), is characterized by low Re concentrations
despite evidence that suggests these sediments are more
reducing (e.g., lower I/Corg ratios, porewater sulphate
depletion, relatively abundant AVS and pyrite).
In marine sediments, two processes can result in U
concentrations above the lithogenic background value
(i.e., ∼1 μg/g for Vancouver Island; Morford et al., 2001).
Uranium can be scavenged from the water column and
delivered to the sediment as part of the particulate flux
(i.e., particulate non-lithogenic U; Anderson, 1982;
Zheng et al., 2002). The exact mechanism responsible
for U scavenging is not yet known. However, it is clear
that the preservation of particulate non-lithogenic U
(PNU) depends on oxygen availability such that PNU is
well preserved in oxygen-poor environments (b 25 μM)
but poorly preserved in well-oxygenated environments
(i.e., Zheng et al., 2002). The second mechanism of U
accumulation is authigenic enrichment under conditions
similar to those required for authigenic Re accumulation
(i.e., suboxic conditions; Cochran et al., 1986; Klinkhammer and Palmer, 1991). In our cores, U is not enriched in
the upper 10 to 15 cm. These results suggest that any PNU
that is delivered to the sediment is not preserved. Furthermore, the weakly suboxic conditions in these surface
sediments are insufficient for authigenic accumulation.
However, there is a relatively large authigenic accumulation (up to 5.8 μg/g) in the lower part of Core 09, and
smaller enrichments in cores 02 and 05 (i.e., the same
cores in which authigenic Re accumulates). These results
are expected given that U, like Re, accumulates under
strongly suboxic conditions.
Cadmium may be enriched by a few μg/g in suboxic
sediments (McCorkle and Klinkhammer, 1991; Rosenthal
et al., 1995b; van Geen et al., 1995) and is often highly
enriched (N 8 μg/g) in anoxic sediments (Pedersen et al.,
1989; van Geen et al., 1995; Morford et al., 2001). The
most probable enrichment mechanism is the precipitation
of CdS (Rosenthal et al., 1995a). Thus, although the
oxidation state of Cd is not affected by redox conditions
(i.e., Cd has a single stable oxidation state in aqueous
solution), its sedimentary enrichment requires the presence of dissolved sulfide. In the oxic and weakly
suboxic near-surface sediments of the Vancouver Island
continental margin, the concentration of Cd is low
(≤0.3 μg/g). This observation is consistent with other
geochemical data that indicate, with the exception of
Core 01, little accumulation of sulfide minerals (Fig. 4).
There are, however, small Cd enrichments in the lower
portion of cores 09, 02 and 05 that suggest minor sulfide
formation is occurring at depth. Once again trace metal
enrichment is greatest in Core 09 (up to 1 μg/g) implying
stronger reducing conditions at this OMZ site. Interestingly, there is no Cd enrichment in Core 01 despite clear
evidence that sulphate reduction and Fe sulfide precipitation are occurring. This apparent contradiction most
probably reflects a combination of two factors. First,
unlike Re, Mo and U that are conservative in seawater,
Cd is a biologically active element characterized by low
concentrations in ocean surface waters, increasing
concentrations down to a water depth of approximately
1000 m, and consistently high concentrations below this
depth. Thus, the concentration of Cd in the bottom
water at Station 01 (120 m water depth) should be
substantially lower than in the bottom waters that overlie
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
193
Fig. 8. Downcore concentration profiles of acid volatile sulfide (AVS), pyrite and Mo in cores 01, 06, and 09 (figures a, b, and c respectively). Note
that pyrite and Mo show similar concentration profiles, consistent with the current model of Mo enrichment in reducing sediments (i.e., adsorption of
Mo onto pyrite). The dashed line in Fig. 8b indicates the contact between the Holocene green mud and Pleistocene gray clay in Core 06.
the deeper stations, in particular Station 09. In other
words, the supply of Cd at Station 01 may be more
limited. Second, the much higher sedimentation rate at
Station 01 will limit diffusion of Cd into the sediment.
There is, however, sufficient Fe delivered to the
sediment with the detrital flux to support the formation
of Fe sulfide at Station 01.
In oxic sediments, Mo concentrations can be high
(N60 μg/g) as a result of Mo adsorption onto Mn oxyhydroxides (Bertine and Turekian, 1973), but under
suboxic conditions such oxides dissolve and adsorbed
Mo is released to the porewater. As a result, suboxic
sediments contain only lithogenic concentrations of Mo
(∼ 0.6 μg/g for terrigenous sediments derived from
eastern Vancouver Island; Morford et al., 2001). In
contrast, anoxic sediments may contain as much as
130 μg/g Mo (Bertine and Turekian, 1973; Koide et al.,
1986; Francois, 1988; Emerson and Huested, 1991;
Crusius et al., 1996; Zheng et al., 2000; Adelson et al.,
2001). Molybdenum enrichment is typically associated
with the precipitation of pyrite (Bertine, 1972; HuertaDiaz and Morse, 1992) but is generally not related to the
precipitation of MoS2. Rather, molybdate (MoO4− 2)
diffuses into the sediment from the overlying water
column where, in the presence of N11 μM H2S, it is
readily converted to thiomolybdate (MoS4− 2) which is
scavenged by pyrite (Bertine, 1972; Helz et al., 1996,
Erickson and Helz, 2000). The presence of zero-valent
sulfur speeds up this process, and also aids in the reduction
of Mo and formation of Mo–Fe–S complexes that are
also rapidly scavenged by pyrite (Vorlicek et al., 2004).
The low Mo concentrations (b 1.4 μg/g) observed in
all cores examined in this study suggest that fully anoxic
conditions do not exist in these sediments. This
interpretation is consistent with the limited formation
of sulfide minerals. Nevertheless, all of the cores,
particularly Core 01, exhibit a slight increase in Mo
concentration with depth that closely follows that of
pyrite (Fig. 8). This suggests that limited sulphate
reduction is occurring, most probably within organicrich, micro-reducing environments.
5. Summary
Despite the shallow oxygen penetration depths (i.e., the
oxic layer is b 1 cm thick) there is no significant enrichment
of redox-sensitive trace elements in the upper 10 to 15 cm
of sediments deposited off Vancouver Island. These results
are consistent with the low extractable-Mn concentrations,
but high I/Corg ratios that indicate near-surface sediments
are poised between Mn and I reduction. Weakly suboxic
conditions appear to be maintained by the combination of
relatively low sedimentation rates and/or deep bioturbation
that allow a small, continual influx of oxidants from the
194
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
overlying water column. Below the bioturbated zone, in
cores from within and below the OMZ (i.e., cores overlain
by low O2 bottom waters), suboxic conditions become
intense enough to allow the accumulation of Re, Cd and U.
However, there is no Mo enrichment because near-surface
sediments do not become anoxic to an extent that allows
Mo to be sequestered.
Core 01 represents an enigma. Based on the porewater data (i.e., large ammonium gradient and evidence
of sulphate reduction), low I/Corg ratios, and higher AVS
and pyrite contents, this core appears to have the most
reducing conditions. However, in apparent contradiction
to this interpretation, redox-sensitive trace metal concentrations are low throughout the core. These observations may reflect the non-steady state conditions that
prevail at Station 01. Given the episodic nature of
upwelling, which influences both organic carbon flux to
the sediment and bottom water oxygen concentration, we
expect sedimentary redox conditions to fluctuate
throughout the year. In contrast, at deeper locations
where bottom water oxygen concentrations are consistently low (e.g., Station 09), sedimentary redox conditions are more stable and favour the accumulation of
redox-sensitive trace metals below the bioturbated zone.
Sedimentation rate and bioturbation depth are the primary factors controlling sedimentary redox conditions on
the Vancouver Island continental margin. The influence of
bottom water oxygen concentration is less important, but
still evident. For example, the highest concentrations of
redox-sensitive trace metals are found in Core 09 collected
from within the OMZ. Unfortunately, in this core, and in
all cores that exhibit an enrichment of redox-sensitive
trace metals, the enrichment occurs well below the sediment–water interface, thus overprinting the earlier trace
metal record. Therefore, in environments characterized by
relatively low sedimentation rates and deep bioturbation,
the paleo-signal derived from redox-sensitive trace metals
may be compromised.
Acknowledgements
We would like to thank P. Anschultz, M. Gehlen and B.
Sundby who provided some of the data used in this paper.
We are also grateful to K. Gordon, C. Guignard, B.
Mueller, and M. Soon for their assistance in the laboratory.
The manuscript has benefited from the reviews of Y.
Zheng and one anonymous referee, as well as from the
editorial handling of D. Rickard. Funding for this project
was provided by NSERC through the Canadian Joint
Global Ocean Flux Study (CJGOFS) and by the CFCAS
through the Climate System History and Dynamics
program (CSHD).
Appendix A. Supplementary data
Supplementary data associated with this article
can be found, in the online version, at doi:10.1016/j.
chemgeo.2006.11.008.
References
Adelson, J.M., Helz, G.R., Miller, C.V., 2001. Reconstructing the rise
of recent coastal anoxia molybdenum in Chesapeake Bay
sediments. Geochimica et Cosmochimica Acta 65, 237–252.
Anabar, A.D., Creasar, R.A., Papanastassiou, D.A., Wasserburg, G.J.,
1992. Rhenium in seawater: confirmation of generally conservative
behaviour. Geochimica et Cosmochimica Acta 56, 4099–4103.
Anderson, R.F., 1982. Concentrations, vertical flux, and remineralization of particulate uranium in seawater. Geochimica et Cosmochimica Acta 46, 1293–1299.
Antoine, D., Andre, J.M., Morel, A., 1996. Oceanic primary production.
2. Estimation at global scale from satellite (coastal zone colour
scanner) chlorophyll. Global Biogeochemical Cycles 10, 57–69.
Bertine, K.K., 1972. The deposition of molybdenum in anoxic waters.
Marine Chemistry 1, 43–53.
Bertine, K.K., Turekian, K.K., 1973. Molybdenum in marine deposits.
Geochimica et Cosmochimica Acta 37, 1415–1434.
Bornhold, B.D., Barrie, J.V., 1991. Surficial sediments on the Western
Canadian Continental Shelf. Continental Shelf Research 11, 685–699.
Bornhold, B.D., Yorath, C.J., 1984. Surficial geology of the continental
shelf, northwestern Vancouver Island. Marine Geology 57, 89–112.
Boyko, T.F., Baturin, G.N., Miller, A.D., 1986. Rhenium in recent
ocean sediments. Geochemistry International 23, 38–47.
Brendel, P.J., Luther III, G.W., 1995. Development of a gold amalgam
voltammetric microelectrode for the determination of dissolved Fe,
Mn, O2, and S(-II) in porewaters of marine and freshwater
sediments. Environmental Science and Technology 29, 751–761.
Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel
in the Black Sea. In: Ittekkot, V., Kempe, S., Michaelis, W., Spitzy,
A. (Eds.), Facets of Modern Biogeochemistry. Springer-Verlag,
Berlin, pp. 326–352.
Carpenter, R., Bennett, J.T., Peterson, M.L., 1981. 210Pb activities in
and fluxes to sediments of the Washington continental slope and
shelf. Geochimica et Cosmochimica Acta 45, 1155–1172.
Cochran, J.K., Carey, A.E., Scholkovitz, E.R., Surprenant, L.D., 1986.
The geochemistry of uranium and thorium in coastal marine
sediments and sediment pore waters. Geochimica et Cosmochimica
Acta 50, 663–680.
Colodner, D., Sachs, J., Ravizza, G., Turekian, K., Edmond, J., Boyle,
E., 1993. The geochemical cycle of rhenium: a reconnaissance.
Earth and Planetary Science Letters 117, 205–221.
Crusius, J., Calvert, S., Pedersen, T., Sage, D., 1996. Rhenium and
molybdenum enrichments in sediments as indicators of oxic,
suboxic and sulfidic conditions of deposition. Earth and Planetary
Science Letters 145, 65–78.
Crusius, J., Pedersen, T.F., Calvert, S.E., Cowie, G.L., Oba, T., 1999.
A 36 kyr geochemical record from the Sea of Japan of organic
matter flux variations and changes in intermediate water oxygen
concentrations. Paleoceanography 14, 248–259.
Dean, W.E., Gardner, J.V., Piper, D.Z., 1997. Inorganic geochemical
indicators of glacial–interglacial changes in productivity and anoxia
on the California continental margin. Geochimica et Cosmochimica
Acta 61, 4507–4518.
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Dean, W.E., Piper, D.Z., Peterson, L.C., 1999. Molybdenum acccumulation in Cariaco basin sediments over the past 24 k.y.: a record of
water-column anoxia and climate. Geology 27, 210–507.
Dionex, 1986. Method for the Determination of Trace Sulfate in Brine.
Application Note, vol. 53. Dionex Corp., Sunnyvale, California.
Eakins, J.D., Morrison, R.T., 1978. A new procedure for the determination of lead-210 in lake and marine sediments. International
Journal of Applied Radiation and Isotopes 29, 531–536.
Edenborn, H.M., Mucci, A., Belzile, N., Lebel, J., Silverberg, N.,
Sundby, B., 1986. A glove box for the fine-scale subsampling of
sediments box-cores. Sedimentology 33, 147–150.
Emerson, S.R., Huested, S.S., 1991. Ocean anoxia and the concentration
of molybdenum and vanadium in seawater. Marine Chemistry 34,
177–196.
Erickson, B.E., Helz, G.R., 2000. Molybdenum(VI) speciation in
sulfidic waters: stability and lability of thiomolybdates. Geochimica et Cosmochimica Acta 64, 1149–1158.
Francois, R., 1987. The influence of humic substances on the
geochemistry of iodine in nearshore and hemipelagic marine
sediments. Geochimica et Cosmochimica Acta 51, 2417–2427.
Francois, R., 1988. A study on the regulation of the concentrations of some
trace metals (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet
sediments, British Columbia, Canada. Marine Geology 83, 285–308.
Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Leudtke, N., Heath,
G.R., Cullen, D., Dauphin, P., Hammond, D., Hartman, B.,
Maynard, V., 1979. Early oxidation of organic matter in pelagic
sediments of the eastern equatorial Atlantic: suboxic diagenesis.
Geochimica et Cosmochimica Acta 43, 1075–1090.
Gagnon, C., Mucci, A., Pelletier, E., 1995. Anomalous accumulation of
acid-volatile sulfides in a coastal marine sediment (Saguenay Fjord,
Canada). Geochimica et Cosmochimica Acta 59, 2663–2675.
Hall, P.O.J., Aller, R.C., 1992. Rapid, small volume flow injection
analysis of CO2 and NH+4 in marine and freshwaters. Limnology and
Oceanography 37, 11119–11313.
Helz, G.R., Miller, C.V., Charnock, J.M., Mosselmans, J.F.W., Pattrick,
R.A.D., Garner, C.D., Vaughan, D.J., 1996. Mechanism of
molybdenum removal from the sea and its concentration in black
shales: EXAFS evidence. Geochimica et Cosmochimica Acta 60,
3631–3642.
Hsieh, Y.P., Yang, C.H., 1989. Diffusion methods for the determination of reduced inorganic sulfur species in sediments. Limnology
and Oceanography 34, 1126–1130.
Huerta-Diaz, M.A., Morse, J.W., 1992. Pyritization of trace metals in
anoxic sediments. Geochimica et Cosmochimica Acta 56, 2681–2702.
Huyer, A., 1983. Coastal upwelling in the California Current System.
Progress in Oceanography 12, 259–284.
Ivanochko, T., Pedersen, T.F., 2004. Determining the influences of late
Quaternary ventilation and productivity variations on Santa
Barbara Basin sedimentary oxygenation: a multi-proxy approach.
Quaternary Science Reviews 23, 467–480.
Klinkhammer, G.P., Palmer, M.R., 1991. Uranium in the oceans: where
it goes and why. Geochimica et Cosmochimica Acta 55, 1799–1806.
Koide, M., Hodge, V.F., Yang, J.S., Stallard, M., Goldberg, E.G.,
Calhoun, J., Bertine, K.K., 1986. Some comparative marine chemistries of rhenium, gold, silver and molybdenum. Applied Geochemistry 1, 705–714.
Kostka, J.E., Luther III, G.W., 1994. Partitioning and speciation of
solid phase iron in saltmarsh sediments. Geochimica et Cosmochimica Acta 58, 1701–1710.
Lord III, C.J., 1982. A selective and precise method for pyrite
determination in sedimentary materials. Journal of Sedimentary
Petrology 52 (2), 664–666.
195
Li, Y.-H., Gregory, S., 1974. Diffusion of ions in sea water and deepsea sediments. Geochimica et Cosmochimica Acta 38, 703–714.
Lucotte, M., d'Anglejan, B., 1985. A comparison of several methods
for the determination of iron hydroxides and associated orthophosphates in estuarine particulate matter. Chemical Geology 48,
257–264.
Luther, G.W., Brendel, P.J., Lewis, B.L., Sundby, B., Lefrancois, L.,
Silverberg, N., Nuzzio, D.B., 1998. Simultaneous measurement of
O-2, Mn, Fe, I-, and S(-II) in marine pore waters with a solid-state
voltammetric microelectrode. Limnology and Oceanography 43,
325–333.
Mackas, D.L., Denman, K.L., Bennett, A.F., 1987. Least squares
multiple tracer analysis of water mass composition. Journal of
Geophysical Research 31, 2907–2918.
Malcolm, S.J., Price, N.B., 1984. The behaviour of iodine and bromine
in estuarine surface sediments. Marine Chemisty 15, 263–271.
Manheim, F.T., 1970. The diffusion of ions in unconsolidated
sediments. Earth and Planetary Science Letters 9, 307–309.
McCorkle, D.C., Klinkhammer, G.P., 1991. Porewater cadmium
geochemistry and the porewater cadmium: (13C relationship).
Geochimica et Cosmochimica Acta 55, 161–168.
McKay, J.L., 2003. Palaeoceanography of the Northeastern Pacific
ocean off Vancouver Island, Canada, PhD Thesis, University of
British Columbia, 237p.
McKay, J.L., Pedersen, T.F., Kienast, S.S., 2004. Organic carbon
accumulation over the last 16 kyr off Vancouver Island, Canada:
evidence for increased marine productivity during the deglacial.
Quaternary Science Reviews 23, 261–281.
McKay, J.L., Pedersen, T.F., Southon, J., 2005. Intensification of the
oxygen minimum zone in the Northeast Pacific during the last
deglaciation: ventilation and/or export production? Paleoceanography 20. doi:10.1029/2003PA000979.
Middelburg, J.J., Soetaert, K., Herman, P.M.J., 1991. Empirical
relationships for use in global diagenetic models. Deep-Sea
Research Part 1 44, 327–344.
Morford, J.L., Emerson, S., 1999. The geochemistry of redox sensitive
trace metals in sediments. Geochimica et Cosmochimica Acta 63,
1735–1750.
Morford, J.L., Russell, A.D., Emerson, S., 2001. Trace metal evidence
for changes in the redox environment associated with the transition
from terrigenous clay to diatomaceous sediment, Saanich Inlet, B.C.
Marine Geology 174, 355–369.
Morford, J.L., Emerson, S.R., Breckel, E.J., Kim, S.H., 2005.
Diagenesis of oxyanions (V, U, Re, and Mo) in pore waters and
sediments from a continental margin. Geochimica et Cosmochimica Acta 69, 5021–5032.
Mortlock, R.A., Froelich, P.N., 1989. A simple method for the rapid
determination of biogenic opal in pelagic marine sediments. DeepSea Research 36, 1415–1426.
Nameroff, T.J., Balistrieri, L.S., Murray, J.W., 2002. Suboxic trace
metal geochemistry in the eastern tropical North Pacific.
Geochimica et Cosmochimica Acta 66, 1139–1158.
Nameroff, T.J., Calvert, S.E., Murray, J.W., 2004. Glacial–interglacial
variability in the eastern tropical North Pacific oxygen minimum
zone recorded by redox-sensitive trace metals. Paleoceanography
19. doi:10.1029/2003PA000912.
Pailler, D., Bard, E., Rostek, F., Zheng, Y., Mortlock, R., van Geen, A.,
2002. Burial of redox-sensitive metals and organic matter in the
equatorial Indian Ocean linked to precession. Geochimica et
Cosmochimica Acta 66, 849–865.
Pedersen, T.F., Waters, R.D., MacDonald, R.W., 1989. On the natural
enrichment of cadmium and molybdenum in the sediments of
196
J.L. McKay et al. / Chemical Geology 238 (2007) 180–196
Ucluelet Inlet, British Columbia. The Science of the Total Environment 79, 125–139.
Price, N.B., Calvert, S.E., 1973. The geochemistry of iodine in
oxidized and reduced recent marine sediments. Geochimica et
Cosmochimica Acta 37, 2149–2158.
Price, N.B., Calvert, S.E., 1977. The contrasting geochemical behaviours
of iodine and bromine in recent sediments from the Namibian shelf.
Geochimica et Cosmochimica Acta 41, 1769–1775.
Ravizza, G., Turekian, K.K., Hay, B.J., 1991. The geochemistry of
rhenium and osmium in recent sediments from the Black Sea.
Geochimica et Cosmochimica Acta 55, 3741–3752.
Reeburgh, W.S., 1967. An improved interstitial water sampler.
Limnology and Oceanography 12, 163–165.
Rosenthal, Y., Boyle, E.A., Labeyrie, L., Oppo, D., 1995a. Glacial
enrichments of authigenic Cd and U in subantarctic sediments: a
climatic control on the elements oceanic budget? Paleoceanography 10, 395–413.
Rosenthal, Y., Lam, P., Boyle, E.A., Thomson, J., 1995b. Authigenic
cadmium enrichments in suboxic sediments: precipitation and
postdepositional mobility. Earth and Planetary Science Letters 132,
99–111.
Strickland, J.D.H., Parsons, T.R., 1972. A Practical Handbook of
Seawater Analysis. Fisheries Research Board of Canada, Bulletin
167, 49–52.
Stuiver, M., Reimer, P.J., Bard, E., Beck, J.W., Burr, G.S., Hughen, K.A.,
Dromer, B., McCormac, G., Van der Plicht, J., Spurk, M., 1998.
INTCAL98 radiocarbon age calibration, 24,000-0 cal BP. Radiocarbon 40, 1041–1083.
Sundby, B., Martinez, P., Gobeil, C., 2004. Comparative geochemistry of
cadmium, rhenium, uranium, and molybdenum in continental margin
sediments. Geochimica et Cosmochimica Acta 68, 2485–2493.
Thomson, J., Higgs, N.C., Croudace, I.W., Colley, S., Hydes, D.J., 1993.
Redox zonation of elements at an oxic-post-oxic boundary in deepsea sediments. Geochimica et Cosmochimica Acta 57, 579–595.
van Geen, A., McCorkle, D.C., Klinkhammer, G.P., 1995. Sensitivity
of the phosphate–cadmium–carbon isotope relation in the ocean to
cadmium removal by suboxic sediments. Paleoceanography 10,
159–169.
Vorlicek, T.P., Kahn, M.D., Kasuya, Y., Helz, G.R., 2004. Capture of
molybdeunum in pyrite-forming sediments: role of ligand-induced
reduction by polysulfides. Geochimica et Cosmochimica Acta 68,
547–556.
Yarincik, K.M., Murray, R.W., Lyons, T.W., Peterson, L.C., Haug, G.H.,
2000. Oxygenation history of bottom waters in the Cariaco Basin,
Venezuela, over the past 578,000 years: results from redox-sensitive
metals (Mo, V, Mn, and Fe). Paleoceanography 15, 593–604.
Zheng, Y., van Geen, A., Anderson, R.F., 2000. Intensification of the
northeast Pacific oxygen minimum zone during the Bolling–
Allerod warm period. Paleoceanography 15, 528–536.
Zheng, Y., Anderson, R.F., van Geen, A., Fleisher, M.Q., 2002.
Preservation of particulate non-lithogenic uranium in marine
sediments. Geochimica et Cosmochimica Acta 66, 3085–3092.