Introduction
Oxygen minimum zones (OMZs) are epipelagic and mesopelagic subsurface layers
of suboxic waters (e.g., ≤ 22 µM O2) found along eastern
boundary currents, the Arabian Sea and the equatorial Pacific, where upwelling of
nutrient-rich waters promotes elevated primary production and O2
consumption through microbial respiration (Wyrtki, 1962; Helly and Levin,
2004; Paulmier and Ruiz-Pino, 2009). Due to strong redox gradients and
reducing conditions, an active microbial community connects cycling of
carbon, nitrogen, sulfur and other elements (Lam et al., 2009; Canfield et
al., 2010; Naqvi et al., 2010; Ulloa et al., 2012; Wright et al., 2012).
Waters overlying the continental shelf of central-southern Chile become
seasonally depleted in O2 during austral spring and summer, when the
area is fed by the poorly oxygenated Peru–Chile Countercurrent. In austral
autumn and winter shelf waters are oxygenated due to the input of
subantarctic waters (Ahumada and Chuecas, 1979; Sobarzo et al., 2007).
Interannual phenomena such as the El Niño–Southern Oscillation (ENSO) can also
affect oxygenation of South Pacific waters (Blanco et al., 2002; Carr et
al., 2002; Levin et al., 2002). In central-southern Chile, the upper edge of
the OMZ deepens during El Niño, thus allowing greater oxygenation of bottom
waters (Gutiérrez et al., 2000; Neira et al., 2001; Escribano et al.,
2004). Analyzing a sedimentary record from northern Chile, Vargas et al. (2007) related changes in coastal upwelling and biological production to
variations in the Pacific Decadal Oscillation (PDO), characterized by an
ENSO-like interdecadal variability in the Humboldt Current System. During
the cool phase of PDO, primary production intensifies in response to
upwelling and fertilization of the upper ocean (Mantua et al., 1997, 2002;
Cloern et al., 2007), leading to enhanced O2 consumption in the water
column (Wyrtky, 1962; Sarmiento et al., 1998; Helly and Levin, 2004). Since
patterns of biological production and oxygenation of the water column during
PDO cycles resemble those of ENSO (Vargas et al., 2007), we hypothesize that
variations at the scale of PDO promote chemical and biological changes in
the OMZ off central-southern Chile.
Trace metals as redox proxies
Past redox variations can be analyzed using trace elements in sediments
since some redox-sensitive metals are less soluble under reducing conditions,
resulting in authigenic enrichment in low-oxygen and high-organic-matter
environments (Algeo and Maynard, 2004; McManus et al., 2005). This chemical
behavior makes molybdenum (Mo), uranium (U), and cadmium (Cd) valuable
paleoredox and paleoproductivity proxies (Calvert and Pedersen, 1993;
Morford and Emerson, 1999; Crusius et al., 1996; Algeo and Maynard, 2004).
Mo occurs primarily as soluble MoO42- in oxygenated marine waters,
and its reduction to particle reactive thiomolybdates
(MoOxS4x2-) under anoxia or molybdenum sulfide
(MoS42-) under euxinia results in authigenic enrichment of
sedimentary Mo (Crusius et al., 1996; Helz et al., 1996; Zheng et al., 2000;
Vorlicek and Helz, 2002), thus indicative of O2-depleted environments.
Uranium is mainly present as U (VI) that binds to carbonate ions, forming
UO2(CO3)34-in seawater. Reduction of U (VI) to U (IV)
occurs under suboxic conditions and at similar redox potentials that allow
Fe(III) reduction to Fe(II) (Cochran et al., 1986; Klinkhammer and Palmer,
1991; Chaillou et al., 2002; McManus et al., 2005, 2006). Higher content of
U relative to Mo indicates anoxic depositional conditions (Algeo and
Maynard, 2004; Tribovillard et al., 2006), whereas equal contents of U, V and
Mo indicate euxinic conditions in the overlying water column (Algeo and Maynard,
2004; Tribovillard et al., 2006).
Cadmium is delivered to marine sediment mainly in association with sinking
organic matter (Piper and Perkins, 2004). If sediments are reduced, then Cd
is authigenically enriched, likely as sulfide (Rosenthal et al., 1995;
Gobeil et al., 1997; Morford and Emerson, 1999; Morford et al., 2001).
Lipid biomarkers
In the past decade, a diverse and active microbial community has been
identified in OMZ waters off central and northern Chile (Stevens and Ulloa,
2008; Farías et al., 2009; Quiñones et al., 2009; Canfield et al., 2010;
Molina et al., 2010; Levipan et al., 2012; Srain et al., 2015). Temporal and
compositional variations in this microbial community can be studied by
analyzing their cell membrane lipids (biomarkers) preserved in the
sedimentary record, as demonstrated in other OMZ areas of the ocean
(Schouten et al., 2000a; Arning et al., 2008; Rush et al., 2012).
Lipid biomarkers are organic molecules occurring in recent and geological
materials that have chemical structures that record their biological origin
(Brassell, 1992; Schouten et al., 2000a; Hinrichs et al., 2003; Coolen et
al., 2008; Talbot et al., 2014). Biomarkers are relatively resistant to
degradation, and they can be indicators of a broad group of organisms or of a
specific genus or species and, as such, of their growing environment (Table 1, Brassell et al., 1986; Brocks and Pearson, 2005). Abundant sedimentary
sterols C27Δ5, C28Δ5, C29Δ5 and C30Δ22 (Volkman, 2003) are indicative
of algal primary and export production. The content and composition of
isoprenoidal glycerol dialkyl glycerol tetraethers (GDGTs) are used as
indicators of ammonia oxidation by marine pelagic archaea (De Long et al.,
1998; Schouten et al., 2000b; Turich et al., 2007; Lincoln et al., 2014),
which are capable of nitrifying under low-O2 conditions (Brandhorst,
1959; Carlucci and Strickland, 1968; Ward and Zafiriou, 1988; Ward et al.,
1989; Lipschultz et al., 1990).
Lipid biomarkers used in this study and their paleobiological
interpretation.
Biomarker
Biological and/or environmental interpretation
References
Hopanoids hydrocarbons
C30 hopanes
Diverse bacterial lineages, few eukaryotic species (e.g., some cryptogams, ferns, mosses, lichens, filamentous fungi, protists)
Rohmer et al. (1984)
Extended C31 to C35 hopanes (homohopanes)
Diagnostic for bacteria; its biosynthesis is restricted to facultative anaerobes and strict anaerobes involved in anaerobic methane cycling (Thiel et al., 2003)
Rohmer et al. (1984), Ourisson and Albrecht (1992)
22, 29, 30-trinor-hop-17(21)-ene (C27 trisnorhopene)
Detected in anoxic and euxinic sediments, and during upwelling events, and considered indicator of anaerobic microbial degradation
Grantham et al. (1980), Volkman et al. (1983), Peters and Moldowan (1993), Schouten et al. (2001)
Hopanols
17β, 21β-hopanol (C30)
Diverse bacterial lineages; diagenetic product of hexafunctionalized bacteriohopanepolyols
Rohmer et al. (1984), Venkatesan et al. (1990), Innes et al. (1997, 1998), Talbot et al. (2001), Farrimond et al. (2002)
17β, 21β-homohopanol (C31)
Diverse bacterial lineages; diagenetic product of pentafunctionalized bacteriohopanepolyols
Rohmer et al. (1984), Venkatesan et al. (1990), Innes et al. (1997, 1998), Talbot et al. (2001), Farrimond et al. (2002)
17β, 21β-bishomohopanol (C32)
Diverse bacterial lineages; diagenetic product of bacteriohopanetetrols
Rohmer et al. (1984), Venkatesan et al. (1990), Innes et al. (1997, 1998), Talbot et al. (2001), Farrimond et al. (2002)
Sterols
C27Δ5
Bacillariophyceae, Bangiophyceae, Dinophyceae, marine Eustigmatophyceae, Haptophyceae; indicator of primary production and algal bloom
Volkman (2003)
C29Δ5
Diverse microalgae lineages (Bacillariophyceae, Chlorophyceae, Chrysophyceae, Euglenophyceae, Haptophyceae, Pelagophyceae, Raphidophyceae, Xanthophyceae); indicator of primary production and algal bloom
Volkman (2003)
C30Δ22
Dinophyceae
Volkman (2003)
MAGEs
C16 MAGE to C18MAGE
Fermentative and sulfate-reducing bacteria; biological source does not appear unique; considered indicators of suboxic/anoxic water column and sediments
Langworthy et al. (1983), Langworthy and Pond (1986), Ollivier et al. (1991), Hernández-Sanchez et al. (2014)
GDGTs
GDGT-0 to GDGT-V
Marine archaea (Thaumarchaeota and Euryarchaeota); considered indicators of ammonia oxidation by Thaumarchaeota and archaeal secondary production
DeLong et al. (1998), Schouten et al. (2000b), Turich et al. (2007), Lincoln et al. (2014)
Changes in sedimentary contents of bacterial hopanes and hopanols are
related to variations in bacterial groups (Rohmer et al., 1984; Ourisson and
Albrecht, 1992; Innes et al., 1998; Talbot et al., 2007). Occurrence of
C27 trisnorhopene is favored in anoxic and euxinic environments, and
during upwelling events (Grantham et al., 1980; Schouten et al., 2001), and
is considered as an indicator of anaerobic microbial degradation (Volkman et
al., 1983; Duan et al., 1996; Duan, 2000; Peters et al., 2005). C16,
C17, and C18 mono-O-alkyl glycerol ethers (MAGEs) are present in
fermentative and sulfate-reducing bacteria (Langworthy et al., 1983;
Langworthy and Pond, 1986; Ollivier et al., 1991), although these biological
sources do not appear to be unique (Hernandez-Sanchez et al., 2014).
We studied redox-sensitive metals and organic biomarkers in a ca. 110-year
sedimentary record from the OMZ within the upwelling ecosystem off
Concepción, central-southern Chile (36∘ S), to infer temporal
changes in biological production and oxygenation of the water column. Our
goal was to assess whether the intensity of the OMZ has varied over the past
century in response to ocean–atmosphere circulation patterns, and whether
this is reflected in changes in the microbial community.
Location of sampling site Station 18 in the upwelling ecosystem
off Concepción, central-southern Chile. Bathymetry in shades of blue,
scale on right-hand side.
Methods
Sampling
The study site (Station 18; 36∘30.8′ S 73∘7′ W) is located in the coastal upwelling ecosystem off central-southern
Chile, ca. 18 nautical miles from the coast of Concepción (Fig. 1).
Sampling was carried out as part of the “Microbial Initiative in Low Oxygen
off Concepción and Oregon” (http://mi_loco.coas.oregonstate.edu) and the Oceanographic Time Series Program
(Station 18) of the Center for Oceanographic Research in the eastern South
Pacific (COPAS) at the University of Concepción
(www.copas.udec.cl/eng/research/serie).
A 25 cm sediment core was collected at a water depth of 88 m during austral
summer (February 2009) using a GOMEX box corer onboard R/V Kay-Kay II. The
top 5 cm was sectioned onboard every 0.5 cm, whereas the rest of the core
was sampled at 1 cm resolution. Samples were stored in glass petri plates and
kept frozen at -18 ∘C until laboratory analysis. The water
column was sampled monthly at Station 18 from January 2008 to November 2009
with Niskin bottles, and temperature, salinity, O2, and fluorescence of
chlorophyll a data were obtained using a Seabird 25 CTDO. Fluorescence data
were transformed to concentration of chlorophyll a according to Parsons et al. (1984). All water column data were obtained from the database of the
COPAS center.
Sedimentary redox potential and organic carbon content
Redox potential was measured in the top 15 cm of the sediment core using a
redox potential sensor (Hanna) with an accuracy of ±0.1 mV.
Sedimentary organic carbon content was determined by high-temperature
oxidation using a NA 1500 Carlo Erba elemental analyzer. Prior to organic
carbon analysis, inorganic carbon was removed by placing samples into silver
cups with a drop of Milli-Q water and then fuming over night with
concentrated HCl. Samples were dried at 60 ∘C for analysis.
Geochronology
Sedimentary 210Pb activities were determined by Alpha spectrometry of
its daughter 210Po using 209Po as a yield tracer (Flynn, 1968).
Activities were quantified until 1σ error was achieved in a Canberra
Quad Alpha Spectrometer. Ages (CE, Common Era) were established according to
the Constant Rate of Supply model (CRS; Appleby and Oldfield, 1978), which
considers unsupported 210Pb inventories (210Pbxs). Geochronology
of our sediment core was determined through radiocarbon measurements on fish
scales and the best fit age curves resulting from CRS model and three
14C control points from a longer core (VG06-2) retrieved in 2006 from
the same sampling site (Muñoz et al., 2012, and Supplement). Resulting
ages were converted to calendar years before present using calibration curve
MARINE09 (Reimer et al., 2009) and applying a regional marine reservoir
correction (ΔR) of 137 ± 164 years, with a 2σconfidence
interval (Table S1 in the Supplement; Fig. 2).
Trace metal analysis
Trace metals Mo, U, and Cd were analyzed with an Agilent 7500ce inductively
coupled plasma mass spectrometer (ICP-MS), and aluminum (Al) was determined
in a Perkin Elmer AAnalyst 700 atomic absorption spectrometer. Sediment
samples and analytical blanks (18.0 MΩ deionized water) were
sequentially digested with Suprapur
HNO3, HCl, HClO4, and HF.
Accuracy and precision of measurements were assessed by analyzing reference
material MESS-3 from the National Research Council of Canada. Excess metal
(Mexs) was calculated as [Mesample] - ((Me/Al)earth × [Alsample]). (Me/Al)earth corresponds to an average
ratio for the Biobío and Itata rivers (Fig. 1) in central-southern Chile
(Muratli, J. M., personal communication, 2012, Table S2).
Geochronology estimated from 210Pbxs inventories (black
line) and 14C measurements ± standard deviation. Ages are years
before present (1950). Dotted line shows fitted values from curve (r2 = 0.99).
Gas chromatography–mass spectrometry (GC-MS) of biomarkers
Extraction of lipid biomarkers (i.e., hopanes, hopanols, sterols, and
MAGEs) from sediments was carried out
according to a modified Bligh and Dyer (1959) procedure, substituting
dichloromethane for chloroform. Freeze-dried sediment samples (1–5 g) were
sequentially extracted by ultra-sonication with 30 mL
dichloromethane / methanol (1:3 v/v, 2×), (1:1 v/v, 1×), and dichloromethane
(2×). The lipid extract was concentrated with a rotary evaporator and dried
with anhydrous Na2SO4. Lipid extracts were then separated into
four fractions by column chromatography (30 cm length, 1 cm ID) filled with
ca. 7 g deactivated silica gel. Aliphatic hydrocarbons (F1) were eluted with
40 mL hexane, ketones (F2) were eluted with 50 mL toluene / hexane (1:3 v/v),
alcohols (F3) were eluted with 50 mL ethyl-acetate / hexane (1:9 v/v), and polar compounds
(F4) were eluted with 35 mL ethyl-acetate / methanol / hexane (4:4:1 v/v). The
alcohol fraction (F3) was derivatized with 80 µL BSTFA (N,O-
bis(trimethylsilyl) trifluoroacetamide) and 40 µL TMCS
(trimethylchlorosilane) at 70 ∘C for 1 h before analysis.
Samples were analyzed in an Agilent 6890 GC series coupled to an Agilent
5972 MS. Hopanols, sterols and MAGEs were analyzed with a 30 m DB-5 column
(0.5 mm ID, 0.25 µm film thickness) using He as a carrier gas. Oven
temperature program included 60 ∘C (2 min) to 150 ∘C at 15 ∘C min-1, to 320 ∘C (held 34.5 min) at 4 ∘C min-1.
Hopanes were analyzed in the aliphatic hydrocarbon fraction (F1) using a 30 m HP-5 column (0.32 mm ID, 0.25 µm film thickness). GC oven
temperature program was 80 ∘C (2 min) to 130 ∘C at 20 ∘C min-1, to 310 ∘C at 4 ∘C min-1. The MS was operated
in electron impact mode (70 eV) with the ion source at 250 ∘C. Mass spectra were acquired in both full scan mode (m/z range 40–600,
scan rate 2.6 s-1) and selective ion-monitoring mode (SIM, m/z 191 for
hopanes and hopanols). Concentrations of alcohols and aliphatic hydrocarbons
were based on those of internal standards 1-nonadecanol and squalene.
Analysis of GDGTs by
high-performance liquid chromatography–atmospheric pressure chemical
ionization–mass spectrometry (HPLC-APCI-MS)
Sedimentary material was sequentially extracted by ultrasonication (3×) with
methanol, dichloromethane / methanol (1:1, v/v), and dichloromethane. Lipid
extracts were concentrated using a rotary evaporator and dried over a small
Pasteur pipette filled with combusted glass wool and anhydrous
Na2SO4. Lipids were separated into non-polar and polar fractions
using a Pasteur pipette filled with activated Al2O3, after elution
with hexane / dichloromethane (9:1, v/v) and dichloromethane / methanol
(1:1 v/v), respectively. An aliquot of the polar fraction was dissolved in
hexane / propanol (99:1 v/v) and filtered through a 0.45 µm PFTE
filter. HPLC-MS analysis followed methodologies described by Hopmans et al. (2000) and Liu et al. (2012), using an Agilent Technologies 1200 Series HPLC
equipped with an auto-sampler and a binary pump, linked to a Q-TOF 6520 mass
spectrometer via an atmospheric pressure chemical ionization interface
(Agilent Technologies). Samples were dissolved in 200 µL
hexane / isopropanol (99:1 v/v). GDGTs were separated using a Prevail Cyano
column (2.1 × 150 mm, 3 mm; Grace, USA) and maintained at 35 ∘C and a flow rate of 0.25 mL min-1. The elution
program was 5 min 100 % eluent A (hexane / isopropanol, 99:1, v/v),
followed by a linear gradient to 100 % eluent B (hexane / isopropanol,
90:10 v/v) for 35 min, and then held at 100 % eluent B for 5 min.
Quantification of core GDGTs was achieved by co-injection of samples with a
C46 GDGT as the internal standard (Huguet et al., 2006).
Statistical analysis
Homogeneity of variances was assessed using Levene's test, whereas
normality was determined using a Shapiro–Wilk test. Non-parametric Spearman
correlations were calculated between selected variables in order to
determine statistical associations with significance < 0.05
(Statistica software, version 12).
Results
Oceanographic setting of the study site
During austral fall and winter (April to August), water temperature ranged
between 11 and 12 ∘C in the upper 20 m of the water column,
and between 10 and 11 ∘C below 65 m depth (Fig. S1a in the Supplement).
Surface salinity varied between 32 and 33 above 20 m, and was 34 below this
depth (Fig. S1b). Chlorophyll a concentration varied between 0.3 and 1.4 mg m-3,
with higher values in the top 20 m (Fig. S1c). Oxygen concentration
varied between 170 and 205 µM in the top 20 m, and dropped to values
lower than 22 µM (suboxia) below 60 m depth (Fig. S1d). During
austral spring and summer (September to March) surface temperature ranged
between 13 and 15 ∘C, decreasing to 10 ∘C
below 84 m depth (Fig. S1a). Salinity varied between 31 and 34.5 in the
whole water column (Fig. S1b). Chlorophyll a concentrations up to 53 mg m-3 were measured in surface waters (Fig. S1c). Oxygen concentration
ranged between 114 and 217 µM in surface waters. Suboxic waters
(i.e., < 22 µM) occur below ca. 20 m (Fig. S1d), which is
significantly shallower than in austral fall–winter.
Redox potential decreased from -176 mV at the water–sediment interface to
-325 mV below 3 cm, indicating predominance of reducing conditions in
near-surface sediments at the time of sampling during austral summer (Fig. S1e), consistent with the occurrence of 5 µM O2 in bottom waters
(Fig. S1d). A surface fluff layer with a Thioploca mat was observed at the
sediment–water interface. Organic carbon content varied between 0.07 and
0.1 g (gdw)- (Fig. S1e).
Geochronology
Background 210Pbxs activity of 0.80 ± 0.02 dpm g-1
was reached at 23 cm in the core. Geochronology from both 210Pbxs
inventories and radiocarbon ages (Fig. 2; Table S1) fitted an exponential
decrease (r2 0.99) due to sediment compaction (Fig. 2), allowing
adjustment of older ages (Binford, 1990). A recent sedimentation rate of 0.24 ± 0.02 cm yr-1 was estimated, representing ca. 110 years of
sedimentation in our sediment core at Station 18.
Redox-sensitive trace metals
Redox-sensitive metals are enriched in the interval ca. 1935–1971 CE (Fig. 3a–c; black bar). Moxs content ranged between 2.5 and 6.5 ppm
(Fig. 3a), showing a similar vertical distribution to Uxs (1.1–4.1 ppm; Fig. 3b) and to Cdxs (0.8–1.9 ppm; Fig. 3c). Enrichments of Moxs,
Uxs, and Cdxs exhibited a significant correlation among each other
(Rs: p < 0.05; Table 2; Fig. S2), indicating more reducing
conditions in bottom waters and sediments at this time. In comparison, the
periods 1905–1919 CE and 1979–2005 CE showed lower contents of
redox-sensitive metals (Fig. 3a–c; white bars), pointing to presumably more
oxygenated bottom waters and sediments.
Downcore excess content (ppm) of redox-sensitive metals (a) Mo,
(b) U, and (c) Cd. Shaded area and black bar correspond to a period of ca.
35 years of enhanced authigenic precipitation of redox-sensitive metals
compared to periods of higher oxygenation (white bars) and low authigenic
precipitation. CE: Common Era. Samples for interval 1957–1969 were
lost.
Spearman rank order correlations. Significant values (p < 0.05) are highlighted in bold.
Moxs
Uxs
Cdxs
Sterols
GDGTs
C27
C31
C32
MAGEs
PDO index
trisnorhopane
hopanol
hopanol
Moxs
0.6
0.6
0.2
0.3
0.2
0.6
-0.3
0.6
-0.3
Uxs
0.6
0.6
0.4
0.6
0.5
0.4
-0.5
0.4
-0.3
Cdxs
0.6
0.6
0.1
0.3
0.6
0.4
-0.4
0.5
-0.3
Sterols
0.2
0.4
0.1
0.3
0.4
0.1
-0.3
-0.4
-0.3
GDGTs
0.3
0.6
0.3
0.3
0.3
0.6
-0.4
0.4
-0.2
C27 trisnorhopane
0.2
0.5
0.6
0.4
0.3
0.5
-0.3
0.4
-0.4
C31 hopanol
0.6
0.4
0.4
0.1
0.6
0.5
-0.4
0.4
-0.3
C32 hopanol
-0.3
-0.5
-0.4
-0.3
-0.4
-0.3
-0.4
-0.4
0.3
MAGEs
0.6
0.4
0.5
-0.4
0.4
0.4
0.4
-0.4
-0.2
PDO index
-0.3
-0.3
-0.3
-0.3
-0.2
-0.4
-0.3
0.3
-0.2
Algal sterols
Sterols C27Δ5, C28Δ5, C29Δ5 and C30Δ22 were identified through the
fragmentation pattern of their trimethylsilyl (TMS) derivatives. The
presence of C27Δ5 sterol cholesterol (m/z 458 [M]+)
was confirmed by detection of ions m/z 129, m/z 329 and 368. The
C28Δ5 sterol (m/z 472 [M]+) showed ions of m/z 129,
as well as m/z 343 and m/z 382. The C29Δ5 sterol (m/z 486
[M]+) was identified by prominent ions m/z 357 and 396. Prominent ions
m/z 69, m/z 271, m/z 359 and m/z 500 [M]+ confirmed the presence of
C30Δ22 dinosterol. Sterol contents ranged between 1029
and 12 164 µg (g Corg)-1, with maximum values in surface
sediments (Fig. 4a). Sterols correlated positively with Uxs (Rs:
p < 0.05; Table 2; Fig. S3).
Archaeal GDGTs
GDGTs were identified by their molecular ion and elution pattern: GDGT-0
(1302 15 [M + H]+); GDGT-I (1300 [M + H]+); GDGT-II (1298 [M
+ H]+); GDGT-III (1296 [M + H]+); and GDGT-V and GDGT-V'
(1292 [M + H]+, known as crenarchaeol and crenarchaeol regioisomer).
Content of GDGTs varied between 1094 and 5423 µg (g Corg)-1
(Fig. 4b), with elevated values at the base core and between ca. 1947 and
1975 CE (Fig. 4b). GDGTs and Uxs contents correlated positively
(Rs: p < 0.05; Table 2; Fig. S4).
Downcore contents of (a) sterols; (b) archaeal GDGTs; (c) 17α-22, 29, 30-trinorhopene
(C27 TNH); (d) 17β,21β-homohopanol
(C31 hopanol); (e) 17β, 21β-bishomohopanol (C32 hopanol);
(f) MAGEs; and (g) Pacific Decadal Oscillation (PDO) index
(http://jisao.washington.edu/pdo/PDO.latest). Units are micrograms per gram
dry weight. Shaded area and black bar as in Fig. 3. Gaps in the record
indicate that biomarker content was under detection limit.
Hopanoid composition and abundance
C27 trisnorhopene (22, 29, 30-trinorhop-17,(21)-ene) was identified based
on its molecular ion fragment m/z 368 [M+-2H+] and fragments
m/z 191 and 231, indicating unsaturation in the ring system (Table 3). Three
diploptene isomers were identified according to their mass spectra:
hop-13,18-ene, neohopene, and hop-22,29-ene (Table 3; Fig. S5a). C30 hopene
diploptene was identified based on its molecular ion (m/z 410
[M+]) and diagnostic ions m/z 395, 299 and 191 (Table 3; Fig. S5a). A
homologous series of C31 to C35 hopanes with αβ
configuration were identified through m/z 191 in the hydrocarbon fraction
(Fig. S5a). Homohopanes C31, C33, C34, and C35 were
present as epimers S and R (Fig. 5a, Table 3), whereas C32 hopane
occurred as the single epimer R (Table 3; Fig. S5a). C27 norhopene and
hopanes C30 and C31 were the only compounds with ββ
configuration (Table 3; Fig. S5a). C31 hopane showed the highest
relative abundance in the homohopane homologous series, with S and R
17α, 21β-homohopane as the predominant one, followed by hopanes
C33 and C34 (Fig. S5a).
Molecules identified in m/z 191 mass chromatogram of aliphatic
hydrocarbon and alcohol fractions from shelf sediments off Concepción
(36∘ S).
Hopanoid hydrocarbons
ID
Molecule
Number of
Molecular
carbon atoms
weight
1
17α-22,29,30-trinorhopane
27
370
2
22, 29, 30-trinor-17(21)-ene
27
368
3
17β-22, 29, 30-trinorhopane
27
370
4
17α, 21α-30-norhopane
29
398
5
17β, 21β-norhopene
27
368
6
17β, 21β-hopane
30
412
7
Neohop-13(18)-ene
30
410
8
17α, 21β-hopene
30
410
9
Hop-22(29)-ene
30
410
10
17α, 21β-homohopane (R)
31
426
11
Diploptene
30
410
12
17α, 21β-bishomohopane (R)
32
440
13
17β, 21β-homohopane
31
426
14S/R
17α, 21β-trishomohopane (S-R)
33
454
15S/R
17α, 21β-tetrahomohopane (S-R)
34
468
16S/R
17α, 21β-pentakishomohopane (S-R)
35
482
Hopanols
17
17β, 21β-hopanol
30
500
18
17β, 21β-homohopanol
31
514
19
17β, 21β-bishomohopanol
32
528
20
17β, 21β-trishomohopanol
33
542
17β, 21β-hopanol (C30); 17β, 21β-homohopanol
(C31); 17β, 21β-bishomohopanol
(C32); and 17β, 21β-trishomohopanol (C33) were
identified by the characteristic ion m/z 191 and by their molecular ions
([M]+ m/z 500, m/z 514, m/z 528, and m/z 542; Table 3; Fig. S5b).
Homologue C32 was the most abundant hopanol (Fig. S5b).
C27 trisnorhopene ranged between 0.03 and 1.1 µg (g
Corg)-1. Maximum values occurred between ca. 1935 and 1971 CE
(Fig. 4c), whereas minimum values were observed during intervals 1905–1928
CE and 1980–2005 CE (Fig. 4c). C27 trisnorhopene correlated positively
with Uxs and Cdxs (Rs: p < 0.05; Table 2; Fig. S6).
The profile of C31 hopanol content varied between 1.1 and 3.7 µg
(g Corg)-1, and it reached the highest value during 1935–1971 CE
(Fig. 4d). Positive correlations among C31 hopanol, Moxs, and
Cdxs were observed (Rs: p < 0.05; Table 2; Fig. S7). In
contrast, C32 hopanol anticorrelated with C31 hopanol, Uxs,
and Cdxs (Rs: p < 0.05; Table 2; Fig. 4e, S8).
MAGE indicators of fermentative and
sulfate-reducing bacteria
Mass spectra of MAGEs showed a base peak ion of m/z 205 characteristics of
monoalkyl glycerol–TMS compounds, which corresponds to cleavage between
carbons 1 and 2 of glycerol moiety, and fragment m/z 445
[M+H-CH3]+ due to loss of methyl group. We identified
C16 MAGE with molecular ion m/z 460 [M]+, C17 MAGE with m/z 474 [M]+, and C18 MAGE with m/z 488 [M]+. Content of MAGEs
(sum of C16, C17, and C18 MAGEs) varied between 9 and 628 µg (g Corg-1; Fig. 4f). MAGE content remained low
(50 µg (g Corg)-1) during 1901–1928 CE (Fig. 4f). From ca. 1935 CE,
MAGE contents increased, reaching the highest value in surface sediments
(Fig. 4f). MAGEs correlated positively with Moxs and Cdxs
(Rs: p < 0.05; Table 2; Fig. S9).
Discussion
Patterns of redox depositional conditions, and primary and export
production
We interpret variations in contents of sedimentary redox-sensitive metals as
changes in oxygenation of bottom waters and surface sediments. This
interpretation agrees with previous observations by Böning et al. (2009)
and Muñoz et al. (2012) for the continental shelf off Concepción, as
well as with authigenic enrichments of U and Mo over the Oregon shelf and
Peru upwelling region associated with O2 depletion and increased
primary production (Scholz et al., 2011; Erhardt et al., 2014).
Higher excess amounts of Mo, U, and Cd during the period between 1935 and
1971 (Fig. 3a, b, c) indicate more reduced depositional conditions.
Favorable conditions for Mo and Cd authigenic enrichments are observed in
bottom water and surface sediments during the upwelling season when high
primary production, low water column O2, and severely low redox potential
in surface sediments occur (Figs. S1c, d, and e;
www.copas.udec.cl/eng/research/serie). However, downcore distribution of
these trace metals could also reflect subtle changes in intensity of O2 depletion over the continental shelf off central-southern Chile. Thus,
from ca. 1932 to 1951 CE, an increase in excess Mo and Cd indicates redox
potential favorable to sulfate reduction and HS- production, at least in
bottom waters (euxinic conditions). Since ca. 1957 to 1969 CE, an increase
in excess U content, coincident with a decrease in excess Mo and Cd, could
indicate a transition to anoxia from previous euxinic conditions since U
enrichment begins when the redox potential reaches that for Fe-oxide
reduction (Cochran et al., 1986; Klinkhammer and Palmer, 1991). The observed
temporal variations in the redox potential evidenced by those subtle changes
in trace metals enrichment could result in readjustment of the microbial
community to changing redox potential of the water column. However,
correspondence of these conditions with changes in organic biomarker
patterns is not necessarily detected in sediments since we assume that
sediment diagenesis is constant for organics but not for redox-sensitive
metals once they reach the sediments.
Downcore distribution of inorganic and organic proxies reveals a period of
ca. 35 years between ca. 1935 and 1971 CE (Figs. 3 and 4; black bar) when
values of redox-sensitive metals (Fig. 3), sterols (Fig. 4a), GDGTs (Fig. 4b), C27 trisnorhopene (Fig. 4c), C31 hopanol (Fig. 4d), and MAGEs
(Fig. 4f) were elevated. Taken together, these patterns allow us to infer
that water column O2 was comparatively lower than during the periods
immediately before and after, in association with enhanced primary
production based on the observed increases of sterols and GDGT
concentrations (Fig. 4a). The two periods with relatively more oxygenated
conditions (ca. 1901 and 1919 CE, and ca. 1979 and 2000 CE; Figs. 3 and 4)
are characterized by low metal enrichments (Fig. 3), a lower content of
bacterial biomarkers related to oxygen-depleted conditions – such as C27
trisnorhopene, C31 hopanol, and MAGEs (Table 2; Fig. 4c, d and f) – and
lower organic matter export as evidenced by low contents of sedimentary
sterols (Fig. 4a) and GDGTs (Fig. 4b).
We suggest that for the period 1935–1971 CE algal export production was
elevated, and that this export is responsible for the increase in
phytoplankton sterols (Fig. 4a), which was concurrent with an increase in Cd
(Fig. 3) and GDGTs (Fig. 4b). An enhanced sinking of organic matter leads to
a subsequent increase in the rate of O2 consumption by microbial
degradation, potentially depleting O2 in the water column (Helly and
Levin, 2004; Canfield, 2006) and sediments. Such conditions lead to Mo, U
and Cd enrichment in sediments. Higher GDGT abundance during this time
(Fig. 4b) may reflect a better preservation of archaeal biomarkers favored
by O2 depletion as demonstrated by Schouten et al. (2004) and Zonneveld
et al. (2010). The positive correlations between sterols, GDGTs, and U
enrichments (Table 2) support this conclusion, since U enrichment occurs in
environments with low O2 concentration and/or high organic matter
deposition (Dezileau et al., 2002; Böning et al., 2009; Tribovillard et
al., 2006; Muñoz et al., 2012).
Changes in microbial communities in response to redox variation
Hopanols C31 and C32 are used to analyze changes in the bacterial
community structure because they are the diagenetic products of
bacteriohopanetetrols (BHPs), which in turn can have different bacterial
sources (Talbot et al., 2003). Hopanol content was dominated by C32 hopanol, as found previously in recent sediments (Buchholz et al., 1993;
Innes et al., 1997, 1998; Talbot et al., 2003). C31 hopanol content was
more elevated between ca. 1935 and 1971 CE (Fig. 4d), with peaks at the
beginning and end of the low-O2 period, and exhibited positive
correlation with Moxs and Cdxs (Rs: p < 0.05; Table 2). Content of C32 hopanol (Fig. 4e), a diagenetic product of BHTs
(Innes et al., 1998; Talbot et al., 2003), mostly produced by heterotrophic
aerobic bacteria (Rohmer et al., 1984), displays a common peak with C31
hopanol (Fig. 4d) between 1920 and 1935, and they are decoupled, concurrent
with enrichment of redox-sensitive metals (Fig. 3a, b and c). Observed
changes in abundance and distribution of C31 and C32 hopanols, in
concomitance with past variations of O2 in the water column at the
study site, are consistent with previous findings by Saenz et al. (2011) and
Kharbush et al. (2013). These authors found that the composition and
abundance of BHPs, the biological sources of hopanoids, change with
decreasing O2 in the water column of the Peruvian margin, Arabian Sea,
Cariaco Basin, and in the eastern tropical North Pacific.
Trisnorhopanes are bacterial lipid markers associated with upwelling and
anoxic depositional environments, although their biological sources have not
yet been identified (Schouten et al., 2001; Peters et al., 2005). The
highest C27 trisnorhopene (Fig. 4c) content occurred during the
proposed period of high primary production and O2 depletion
(1935–1971 CE), suggesting a relationship between its abundance and
upwelling-favorable conditions and anaerobic bacterial activity, as
previously suggested for other areas of the world (Grantham et al., 1980;
Duan et al., 1996; Duan, 2000; Schouten et al., 2001).
Sedimentary content of MAGEs was also higher in the period 1935–1971 CE and
in the topmost sediments (Fig. 4f), resembling C16
MAGE (Fig. 5 of
Arning et al., 2008) at the same sampling site (Station 18), assuming
similar sedimentation rate as in our core. MAGEs have been detected in
sediments from upwelling regions of Namibia, Peru, and central-southern
Chile and are attributed to the occurrence of sedimentary sulfate-reducing
bacteria (Arning et al., 2008). The presence of sulfate-reducing bacteria
has been previously documented in coastal waters off Chile (Canfield et al.,
2010) and Peru (Finster and Kjeldsen, 2010).
Forcing mechanisms and variations in OMZ intensity in
central-southern Chile
Combined records of redox-sensitive metals and biomarkers suggest the
occurrence of enhanced reducing conditions at the sediment–water interface
and likely in the water column, from ca. 1935 until 1971 CE (Figs. 3 and 4),
which roughly coincides with a cool (negative) phase of the PDO (Fig. 4g). This suggests a link between changes in
continental shelf oxygenation off Concepción and the PDO, with alternating
phases of decreased (1901–1930 and 1979–1997 CE) and enhanced upwelling
(ca. 1935 to 1971 CE). PDO is a recurring pattern of ocean–atmosphere
variability with phases that last between 2 and 3 decades (Mantua et
al., 1997, 2002). During cool or negative phases, the western Pacific
becomes warmer while parts of the eastern Pacific become colder. The reverse
pattern occurs during warm or positive phase. PDO plays a major role in
decadal-scale oceanographic variability in the Pacific Ocean (Mantua et al.,
1997, 2002; White and Cayan, 1998; Johnson and McPhaden, 1999).
Negative correlations between sedimentary C27 trisnorhopene, C31
hopanol, MAGEs, and PDO values (Table 2; Fig. S10) and a positive
correlation between C32 hopanol (Table 2; Fig. S10) and PDO suggest
that this basin-wide climatic anomaly has an impact on local oceanographic
conditions off Concepción, which in turn modulate the structure of the
microbial community. Bacterial C31 hopanol and MAGEs derive from
microorganisms associated with marked chemoclines and redox gradients
(Rohmer et al., 1984; Innes et al., 1997, 1998; Talbot et al., 2003, 2007;
Kool et al., 2014). Thus, positive PDO phases (warm) were likely associated
with a decrease in wind-driven upwelling, greater oxygenation, decreased
primary productivity, and a concomitant decrease of microorganisms
associated with low O2. Reverse conditions must have dominated during
negative PDO phases, with enhanced upwelling and primary production. An
increase in coastal upwelling off Concepción, as expected during cool
(negative) PDO phases, could contribute to accumulation of atmospheric
greenhouse gases as reported for upwelling ecosystems at seasonal scales
(Bakun and Weeks, 2004; Naqvi et al., 2010).