Occurrence of 1-O-monoalkyl glycerol ether lipids in ocean waters and sediments

Organic Geochemistry 66 (2014) 1–13

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Occurrence of 1-O-monoalkyl glycerol ether lipids in ocean waters and sediments Maria T. Hernandez-Sanchez a,1, William B. Homoky b, Richard D. Pancost a,⇑ a b

Organic Geochemistry Unit, The Cabot Institute and Bristol Biogeochemistry Research Centre, School of Chemistry, University of Bristol, Cantock’s Close, Bristol BS8 1TS, UK Ocean and Earth Science, National Oceanography Centre Southampton, University of Southampton, Southampton SO14 3ZH, UK

a r t i c l e

i n f o

Article history: Received 24 June 2013 Received in revised form 8 September 2013 Accepted 1 October 2013 Available online 10 October 2013

a b s t r a c t The occurrence of 1-O-monoalkyl glycerol ethers (MAGEs) in marine sediments is typically attributed to the presence of sulfate reducers. However, little is known about other possible sources within the marine realm. We have examined the concentrations and distributions of MAGEs in suspended particulate organic matter from surface water and underlying sediments of the Southern Ocean and the eastern South Atlantic Ocean. MAGEs (including monounsaturated, diunsaturated and branched) were found in surface water from both settings, suggesting a planktonic source, likely aerobic bacteria. Supporting this, we found a significant correlation between the concentrations of phytoplankton biomarkers and MAGEs. MAGE distributions in Southern Ocean and South Atlantic surface water differed, with longer chain length and more unsaturated components in Southern Ocean water, which we suggest could be an adaptation to lower surface temperature. Sedimentary MAGE distributions were significantly different from those of overlying water, which we propose to be the result of an additional sedimentary source. However, MAGEs were present in both oxic and nitrogenous–ferruginous sediments, suggesting other bacterial sources besides sulfate reducers. Ó 2013 Elsevier Ltd. All rights reserved.

1. Introduction Glycerol lipids bearing ether moieties are widespread in nature as they are common constituents of cell membranes (including archaeal membranes; Kates, 1993; Koga and Morri, 2005). Vinyl ether lipids [plasmalogens; e.g. 1-O-(alk-10 -enyl)-2-acyl-sn-glycero-3-phospholipids] are ubiquitous in animal cells (Karnovsky et al., 1946; Mangold and Weber, 1987; Nagan and Zoeller, 2001) and widespread in anaerobic microorganisms (Kamio et al., 1969; Fischer et al., 1994) but are apparently rare in plants (Mangold and Weber, 1987). Similarly, alkyl glycerols (generally found as 1-O-alkyl-2,3-diacyl-sn-glycerols and 1-O-alkyl-2-acyl-sn-glycero3-phosholipids) are also widespread among land animals (Karnovsky et al., 1946), although often as minor components, and it has been demonstrated that compounds containing a 1-O-alkyl-snglycerol motif are ubiquitous components of animal cell membranes (Carter et al., 1958). 1-O-alkyl-2,3-diacyl-sn-glycerol lipids have been widely reported in fats of marine animals, including fish tissue (Hallgren et al., 1974), fish oil (Hallgren and Larsson, 1962) and shark liver oil (Wetherbee and Nichols, 2000; Jayasinghe et al., 2003), but also occur in marine invertebrates such as sponges ⇑ Corresponding author. E-mail address: [email protected] (Rich. Pancost). Present address: Geology Department, University of Oviedo, Jesus Arias de Velasco sn, 33005 Oviedo, Spain. 1

0146-6380/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.orggeochem.2013.10.003

(Magnusson and Haraldsson, 2011), molluscs (Hallgren et al., 1974; Clarke, 1984; Hayasaki and Kishimura, 2002; Zhukova, 2007), crustaceans (Hallgren et al., 1974; Fricke et al., 1986) and corals (Imbs et al., 2006). Structurally similar compounds (e.g. 1-O-alkyl-2-acyl-sn-glycerols) have also been found in cultures of sulfate reducing bacteria (SRB; Goossens et al., 1986; Rütters et al., 2001) and marine sediments (Rütters et al., 2002). Related to the above compounds, non-isoprenoid 1-O-monoalkyl glycerol ethers (MAGEs) have been reported in cultured mesophilic SRB (Rütters et al., 2001), as well as in a range of anaerobic fermentative organisms and SRB (Ollivier et al., 1991; Langworthy et al., 1983; Langworthy and Pond, 1986) and sulfur oxidising bacteria (Aquificales genera; Jahnke et al., 2001) isolated from hot spring microbial mats and a hyperthermophilic bacterium isolated from hot marine sediments (Huber et al., 1992). They have also been reported in a range of terrestrial geothermal sediments (e.g. Zeng et al., 1992a,b; Pancost et al., 2005, 2006) and marine sediments subjected to hydrothermal activity (Guaymas Basin), with a new phylogenetically deeply branching bacterial phylotype suggested as the possible source (Teske et al., 2002). In addition, they occur at methane cold seeps associated with the anaerobic oxidation of methane (Hinrich et al., 2002; Pancost et al., 2000; Orphan et al., 2001) and have been suggested to derive from SRB partners of methanotrophic Archaea (Hinrich et al., 2002; Orphan et al., 2001; Elvert et al., 2005). Consequently, the presence of MAGEs in sediments has been linked primarily to the S cycle within

2

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

marine/oceanic environments; for example, MAGEs in sediments from the Namibian, Peru and Chile upwelling regions have been attributed to sedimentary SRB (Arning et al., 2008). Ascribing sedimentary MAGEs to SRB or related sedimentary bacteria, and therefore their potential as biomarkers for sulfate reduction in the geological record, requires that alternative sources can be precluded, but this is challenging given the widespread occurrence of glycerol lipids bearing ether moieties in marine animals (e.g. plasmalogens). Hydrolysis of such compounds – either in the environment or during analysis – would also yield MAGEs and represent an alternative source. Unfortunately, little is known about the occurrence of MAGEs in suspended particulate matter in the oceanic water column. In this study we show that MAGEs are indeed present in suspended particulate organic matter (SPOM) and underlying surface sediments collected from the eastern South Atlantic and the Indian sector of the Southern Ocean (SO). We discuss their distribution and concentration and compare these with algal biomarkers in order to evaluate possible sources under different productivity and geochemical regimes. 2. Methods 2.1. Oceanographic settings SPOM, sinking particles and sediment samples were collected from two different areas: The Crozet Plateau (Indian sector of the

a

b

SO; 45–47°S, 49–51°E; Fig. 1a) and the eastern South Atlantic Ocean (ESAO) off Cape Town (South Africa; Fig. 1b). The area north of the Crozet Islands in the SO experiences an algal bloom during austral summer as a result of natural Fe fertilisation (Planquette et al., 2007; Pollard et al., 2009), whereas the south of the Plateau is more similar to a typical SO high nutrient, low chlorophyll (HNLC) region. SPOM, sinking material and surface sediments were collected from the Crozet Plateau during austral summer 2004 and 2005 from the Fe fertilised area (north) and from the HNLC area (south) during two expeditions (D286 and D300) of the RV Discovery. SPOM was collected from four different stations in the polar frontal zone (PFZ) over two Austral summer seasons (2004/2005 and 2005/2006). During the 2004/2005 campaign, this included stations M3 and M5, within the Fe fertilised area north of the Plateau and M6, within the HNLC region south of the Plateau; the 2005/2006 sampling season included stations M5, M6, and M10, within the Fe fertilised region north of the Plateau (Fig. 1a). Productivity variation during these two austral seasons is described by Hernandez-Sanchez et al. (2010). Sinking particulate material was collected only in the north of the plateau (M10) and surface sediments were collected from north – M5 (4269 m depth) and M10 (3227 m) – and south of the Plateau (M6; 4268 depth) during austral summer 2004/2005 (M5 and M6) and 2005/2006 (M10). The ESAO around 40°S is close to the Sub-Tropical Front (STF) and experiences a range of spatially distinct hydrographic and nutrient conditions. The latitude is characterised by relatively high productivity, especially close to shore. Deep water at this latitude is influenced by Antarctic Bottom Water (AABW), North Atlantic Deep Water (NADW) and Antarctic Intermediate Water (AIW). The coastal area is under the influence of the Aghulas current and a weak upwelling cell is observed north of Cape Town (Shannon, 1985; Lutjeharms and Meeuwis, 1984). Associated with the upwelling cell, a front develops at the shelf break off the Cape Peninsula, which is particularly strong in summer (e.g. strong upwelling and advection of warm Aghulas current water offshore; Shelton and Hutchings, 1990). Associated with the upwelling front, a current flows northwards around the Cape Peninsula (Cape jet; Shelton and Hutchings, 1982). This circulation pattern makes the coastal area off Cape Town a dynamic regime. In contrast, the oceanography of the offshore area is more stable (HardmanMountford et al., 2003) and solely under the influence of the subtropical front. SPOM from surface water (and up to 600 m) and underlying sediments was collected along a transect spanning the shelf– slope–basin transition during the 2010 UK-GEOTRACES expedition of the RV Discovery (D357). SPOM was collected from throughout the water column at 4 offshore stations (St. 1, St. 3, St. 4.5 and St. 6; Fig. 1b) and from the surface (ca. 5 m depth) at three nearshore stations (St. 0, St. 0.5 and St. 0.75); sediments were recovered from all seven stations along this transect (St. 0, St. 0.5, St. 0.75, St. 1, St. 3, St. 4.5 and St. 6). 2.2. SPOM and sinking particles

Fig. 1. Sampling stations around the Crozet Plateau (Southern Ocean; Panel a) and off South Africa (Eastern South Atlantic Ocean; Panel b).

SPOM was collected from the surface mixed layer in SO and ESAO waters. Seawater was filtered in situ using stand alone pumps (SAPs) deployed for 0.5–2 h (ca. 500 l filtered seawater). Additional surface water samples were collected from the ESAO at stations 0, 0.5 and 0.75 from the towed fish (ca. 5 m depth) and filtered on board through GFF filters (293 mm diameter). Samples were handled and stored as described by Hernandez-Sanchez et al. (2012). Sinking particles were collected only at the Crozet Plateau (see Hernandez-Sanchez et al., 2012, for more information). A set of McLane sediment traps (21 cups) was deployed north of the Plateau (M10; 2000 m depth; Fig. 1a) in austral summer 2004–2005

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

and recovered during a second cruise (D300), which took place in December 2005–January 2006, recovering 94% of the annual flux (late December 2004/January 2005–December 2005; Pollard et al., 2009). Sinking particles were handled and stored as described by Hernandez-Sanchez et al. (2012). 2.3. Surface sediments Short mega cores (recovered with a Connelly mega corer, which provided overlying water and the sediment–seawater interface) were collected near the Crozet Plateau at stations M5 (4269 m) and M6 (4268 m) and M10 (3227 m). Similarly, short mega cores were recovered from the South Atlantic UK-GEOTRACES transect at stations 0 (242 m), 3 (4900 m), 4.5 (5269) and 6 (4928 m; Fig. 1b). Sediments from stations 0.5 (733 m), 0.75 (1183 m) and 1 (2602 m) were recovered using a box corer subsequently subsampled using polycarbonate push corers. Cores were cut into 1– 2 cm slices and frozen at 20 °C after sampling. Only the top of the cores (upper 2 cm for SO sediments and upper 10 cm for ESAO sediments) were used. 2.4. Biomarker analysis Freeze dried sediments and GFF filters containing SPOM were hydrolysed for 4 h using 0.1 M KOH in MeOH (5% water) and extracted with a 3:2 (v:v) mixture of dichloromethane (DCM):MeOH in an ultrasonic bath for 15 min (4). The liquid phase was collected and rotary evaporated. Subsequently, for the SO samples, 10 ml double distilled water and 10 ml hexane: DCM (9:1) were added to the total lipid extract (TLE), followed by liquid–liquid extraction of the neutral fraction (4). The same procedure was used for the ESAO samples, except that 10 ml DCM, rather than hexane:DCM, were used for extraction. Freeze dried GFF filters containing sediment trap material were extracted in an ultrasonic bath for 45 min with a 9:1 (v:v) mixture of DCM:MeOH. Half of the TLE was transferred to reaction vials for transmethylation: 1 ml of 30:1 (v:v) MeOH:MeCOCl was added to each sample, which was wrapped in Al foil (to avoid light exposure) and left in a heating block at 45 °C overnight. The extract was then dissolved in DCM and dried by eluting it through a pipette filled with ashed K2CO3. Gas chromatography (GC) and GC–mass spectrometry (GC–MS) were performed as described by Hernandez-Sanchez et al. (2010). Compounds were quantified using 5a-androstan-3b-ol and C37 nalkane (only for alkenones), which had been added to the TLE prior to extraction of the neutral fraction. Compounds extracted from sediment trap material were quantified using 5aH-cholestane as internal standard, which had been added to the filters prior to extraction. 2.5. Porewater oxygen profiling A suite of Unisense micro sensor profiling equipment was used for high resolution downcore O2 determination for the upper 6 cm of ESAO cores. Analysis was performed at sea immediately after recovery of sediments at in situ temperature in a controlled temperature laboratory. Porewater O2 concentration was measured with a Unisense OX100 micro sensor, controlled by a micromanipulator and connected to a Picoammeter and lap top computer by analogue–digital converter. The micro sensor was calibrated before each use by recording the electrical potential difference of an air saturated (oxic) and N2 saturated (anoxic) seawater sample of equivalent temperature and salinity to the overlying bottom water in a Cal 300 chamber. A two point linear interpolation between oxic and anoxic readings was used to determine O2 saturation and converted to molar units with the software SensorTrace Pro

3

using the empirical relationships for O2 saturation in seawater. Micro sensor determinations were made at 200 lm stepped intervals downcore. Repeating the micro-sensor calibration after analysis assessed accuracy, and O2 saturation was reproducible to within 2%. A 50 Hz filter was used to improve signal stability of the micro sensor at sea, where multiple ship operations can introduce minor electrical interference to analysis; Station 6 suffered from an unusual degree of interference, so a 3 point moving average of the data is presented (with a corresponding depth resolution of 600 lm). The O2 penetration depth (OPD) was derived from the depth at which the O2 concentration appeared at or below the micro sensor detection limit (60.3 lM).

3. Results 3.1. Lipid biomarkers A variety of lipids (other than MAGEs), including sterols, alkenones and long chain alkyl diols, are present in SO waters and sediments and have been described by Hernández-Sánchez et al. (2012). Similarly, several sterols, long chain alkenones, phytol and the C28 1,14-diol and C29 12-hydroxymethyl alkanoate, were found in ESAO SPOM and sediments. Cholesterol concentration ranged from 0.19 to 370 ng/l in SPMO (Fig. 2a) and from 16 to 1800 ng/g in underlying sediments (Fig. 2b). Sterols other than cholesterol included 24-methylcholesta-5,22-dien-3b-ol, 24-methylcholesta-5,24(28)-dien-3b-ol and dinosterol among others, with total concentration ranging from 26 to 1200 ng/l in SPOM and from 12 to 3900 ng/g in sediments. The sum of long chain alkenones (including di, tri and tetra unsaturated C37 ketones and diunsaturated C38 ketones), phytol, the C28 1,14-diol and the C29 12hydroxymethyl alkanoate (the latter two only present in sediments) varied from 4.2 to 190 ng/l in SPOM (Fig. 2c) and from 680 and 32,000 ng/g in underlying sediments (Fig 2c). A variety of MAGEs was found in all samples from the ESAO and SO, including SPOM, sinking particles and sediments, and were partially identified using GC–MS. The absolute stereochemistry of the glycerol moiety was not determined; however, since all MAGEs were apparently non-isoprenoid (see below), we tentatively ascribe them to a bacterial origin and sn-1,2 stereochemistry, hence 1-O-monoalkylglycerols. Assignment of the primary saturated components was based on the characteristic mass spectra (Egge, 1983; Zeng et al., 1992a), including an m/z 205 base peak arising from cleavage between C-1 and C-2 of the glycerol and major m/z 117 and 147 fragments, as well as m/z 130 and 133 fragments typical for ether lipids. M+ is not observed under standard EI conditions used, but the presence of [M-15]+, [M-90]+ and [M-103]+ fragments allowed saturated alkyl chain length to be determined (Fig. 3a and b). These assignments were confirmed by the fragment ions associated with the alkyl moiety itself, i.e. m/z 313 for C16, 341 for C18, etc. This revealed that the alkyl chain length varied from C12 to C20 with the C14, C16 and C18 homologues typically dominant (Fig. 4b); the regular retention time spacing of the major homologues, similar to that observed in other settings (e.g. Jahnke et al., 2001; Pancost et al., 2005, 2006), is consistent with these being straight chain rather than branched components. However, branched components, especially branched C15 homologues, assigned on the basis of retention time differences, were present in subordinate abundance in many of the samples. We also tentatively assigned clusters of 4–6 compounds eluting before the C16 and C18 MAGEs as unsaturated MAGEs (Fig. 4b). They had the characteristic m/z 205, 117 and 147 fragments, suggesting that they were structurally related. However, they lacked the strong [M-15]+, [M-90]+ and [M-103]+ fragments from MAGEs and instead had relatively stronger M+ at m/z 458 (Fig. 3c) and

90

r= 0.39

80

p= 0.16

a

70 60 50 40 30 20 10 0 0

20

40

60

80

Phytoplankton biomarkers (ng/g)

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

Cholesterol (ng/L)

4

d

35000 30000 25000 20000 15000 10000 5000 0 0

100 120 140

200

b

r= 0.31 p= 0.06

0

200

400

600

800

Phytoplankton biomarkers (ng/g)

Cholesterol (ng/g)

2000 1800 1600 1400 1200 1000 800 600 400 200 0

4500 3500 3000 2500 2000 1500 1000 500 0

1000

0

200

200 150 100 50 0 0

20

40

60

80

100 120 140

Phytoplankton biomarkers (ng/g)

Phytoplankton biomarkers (ng/L)

p= 0.007

400

600

MAGEs (ng/g)

c

r= 0.66

600

e

r = 0.77 p= 0.004

4000

MAGEs (ng/g) 250

400

MAGEs (ng/g)

MAGEs (ng/L)

35000

f

r = 0.44 p= 0.10

30000 25000 20000 15000

Nitrogenous-ferrigenous ESAO sediments

10000

Oxic ESAO waters and sediments

5000

SO waters 0 0

MAGEs (ng/L)

200

400

600

MAGEs (ng/g)

Fig. 2. (a) Correlation between total MAGEs and cholesterol concentrations in SPOM from ESAO and SO waters. (b) Correlation between total MAGEs and cholesterol concentrations in ESAO sediments. (c) Correlation between total MAGEs and phytoplankton biomarker concentrations (sum of long chain alkenones and phytol) in ESAO and SO SPOM. (d) Correlation between total MAGE and phytoplankton biomarker concentrations (sum of long chain alkenones, phytol, C28 1,14-diol and C29 12-hydroxymethyl alkanoate) in sediments (both oxic and anoxic) collected from the ESAO. (e) Correlation between MAGE and phytoplankton biomarker concentrations in oxic ESAO sediments. (f) Correlation between MAGEs and phytoplankton biomarker concentrations in anoxic ESAO sediments.

486, indicating that they were anhydro C16 and C18 components, respectively. They each eluted prior to the corresponding saturated component, displaying similar chromatographic behaviour to unsaturated fatty acids (FAs) under these GC conditions but being distinct from ring-bearing FAs which eluted later. As such, we tentatively assigned them as a suite of double bond bearing C16 and C18 MAGEs (alkyl chains C16 and C18 respectively; Zeng et al., 1992a). Unfortunately, their low abundance precluded direct testing of this by examining the HI-cleaved hydrocarbons (e.g. Pancost et al., 2001) or assignment of double bond position. 3.1.1. Eastern South Atlantic Ocean Saturated and unsaturated MAGEs ranged from C12 to C20 components. The concentrations are shown in Table 1 and the summed concentration varied from 2.6 to 360 ng/l seawater. The distribution was dominated by the C16 saturated component at most but not all stations (e.g. distribution dominated by C18:0 MAGE at station 0.75), followed by the C14, C17 and C18 saturated components depending on the station (Fig. 5). Branched (C15 MAGEs) and unsaturated (e.g. C16 and C18 MAGEs) components were present in most samples and their concentration (in total) varied from 0.22 and 68 ng/l. In the sediments, branched, unsaturated and saturated MAGEs were also present but the distribution was somewhat more restricted, with only the C15 to C20 components detected. Their

summed concentration varied from 15 to 830 ng/g dried sediment (Table 2) with the C16 (monounsaturated or saturated) or C17 (saturated) MAGEs being dominant (Fig. 6). Like the overlying water, branched C15 MAGES were present in all samples, but branched C17 MAGES were also detected at some stations (e.g. St. 0 and St. 0.5). Unsaturated (e.g. C16 and C18 and sometimes C17) MAGEs were present in most samples and the total concentration (branched and unsaturated) varied from 1.5 ng/g to 210 ng/g. 3.1.2. Crozet Plateau (Southern Ocean) Saturated and unsaturated MAGEs in SPOM from the Southern Ocean water ranged from C14 to C20. In naturally Fe fertilised water (north of the Crozet Plateau), they comprised C14 to C20 components (Fig. 7) and the summed concentration ranged from 4.4 to 1200 ng/l (highest at the highest productivity site M10), with the unsaturated C14 or the diunsaturated C16 MAGEs dominant. Like ESAO water, branched (C15) and unsaturated (C16 and C18) MAGEs were present in most samples (Fig. 7) and ranged (in sum) from 1.3 to 1000 ng/l. Sinking particles collected with sediment traps north of the plateau contained, however, only saturated C16 MAGEs, with a concentration ranging from 0.8–16 ng/g. The underlying sediments contained only the diunsaturated C16 MAGE (Table 3). MAGEs in surface water south of the plateau (in HNLC water) ranged from C15 to C19 components (Fig. 7), with summed

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

5

a

.

b

.

c

.

Fig. 3. Mass spectra of peaks assigned as C16:0 MAGE (a) C14:0 MAGE (b) and C16:1 MAGE in ESAO SPOM.

concentration between 9 and 27 ng/l. Like the north of the plateau, the distributions were dominated by the diunsaturated C16 components, and C15 branched MAGEs were present in all samples; however, C17 branched MAGEs were also present. Their concentration

(branched and unsaturated components) ranged from 6.5 to 23 ng/l. Unfortunately, sinking particle material collected south of the plateau was not available for biomarker analysis. For sediments collected south of the plateau, only the C15 and C16

6

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

a

TIC

Phytol Sterols

St. 6 100 m

IS

MAGES

C16:0 MAGE

Alkenones

m/z 205

C20:0 MAGE

C19:0 MAGE

C18:1 MAGE C18:1 MAGE C18:0 MAGE

C18:2 MAGE

C17:0 MAGE

C16:0 MAGE C16:1 MAGE C16:1 MAGE

C15:0 MAGE

Br C16 MAGE Br C16 MAGE

C14:1 MAGE

C13:0 MAGE

C12:0

C14:0 MAGE

b

Fig. 4. Partial gas chromatograms of neutral fraction of SPOM from the ESAO. (a) Total ion current (TIC) showing major biomarker groups in the samples, including sterols, alkenones and phytol. IS, internal standard (5a-androstan-3b-ol); shaded area represents typical region where MAGEs elute. The area is magnified in panel b (m/z 205), which shows the typical MAGEs distribution (in this particular case the chromatogram corresponds to SPOM collected from the ESAO, Station 6 at 100 m depth), typically comprising C12:0 to C20:0 MAGEs.

(saturated and unsaturated) MAGEs were detected (Table 3) with the branched C15 MAGE dominant.

4. Discussion 4.1. Comparison of water column and sedimentary MAGE distributions and implications for their sources

3.2. Porewater O2 concentration (ESAO sediments) The nearshore ESAO stations (St. 0, St. 0.5, and St. 0.75) and off shore intermediate St. 1 exhibited rapid O2 depletion, with porewater concentration approaching zero from a depth of 0.7– 3.1 cm, with St. 0.75 having the shallowest O2 penetration depth (Fig. 8). Concentration also decreased at the deepest offshore Stations 3, 4 and 6, but remained oxic throughout the upper 6 cm (maximum depth of determinations), with the highest concentrations at station 6.

4.1.1. Sources of MAGEs in the water column MAGEs in marine sediments are typically attributed to SRB (Hinrich et al., 2002; Orphan et al., 2001; Arning et al., 2008). However, at all the sites, the water column is oxic and not characterised by evidence for sulfate reduction, e.g. elevated HS concentration (UK-GEOTRACES consortium, unpublished data). Therefore, it is unlikely that significant sulfate reduction occurs in either the ESAO or SO waters sampled. Sulfate reduction can, however, occur in anoxic microenvironments in an otherwise oxic

7

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13 Table 1 Abundance (ng/l) of MAGEs in SPOM from Eastern South Atlantic Ocean (ESAO) water. St. 3

10 m

10 m

100 m

1.4

0.72

1.10

4.10 1.70 0.56 3.2 7.3

11 0.61

3.1

0.61 2.10

0.41 0.9 4.08 0.64 0.61

1.9

2.4

9.05 0.38 1.3

17.0

2.5

1.5

0.71

16 1.8 4.05

2.6

101

44

3.4

Total

28

12

16

St 6 10 m 50

ng/l

40 30 20 10 0 25

St 0.75 5 m

ng/l

15

10 5

10 m

1.97

12

12

10

10

8

8

6

6

4

4

2

2

0

0 160

St 0.5 5 m

3.7

0.14 0.082 0.15 0.60

11

1.8

22

16

St 4.5 10 m

200 m

2.4 2.76 3.2

14

14

140

12

120

10

100

2.9

St 3 10 m

St. 0.75

St. 0.5

St. 0

5m

5m

5m

4.02 16 0.68 0.48 3.3 0.68 1.9 18 2.4 4.9 24 2.1 2.7

1.6 13 0.64 0.15 0.82 0.64 3.5 14 1.7 8.

19 17 47

83

48

360

8.02 26 140

St 1 20 m

40 30 20 10 0

St 0 5 m

MAGEs

80

6

60

Monounsaturated MAGEs

4

40

Branched MAGEs

2

20

Saturated MAGEs

0

0

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

93 6.09

50

8

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

0

20 m

14

16

20

St. 4.5

3.3 1.1 2.2 1.02 1.5 5.7 52 7.3 2.7

0.39

120

a

200 m

St. 1

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

C12:0 C13:0 C14:0 C15 Br C15 Br C15:0 C16:1 C16:1 C16:0 C17:0 C18:1 C18:1 C18:0 C19:0 C20:0

St. 6

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

MAGEs

MAGEs

b St 3 10 m

4

12

St 3 100 m

3

8

ng/l

St 3 200 m

10

10

6

8 6

2

4

4 1

2

0

0

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

2

0

12: 0 13: 0 14: 0 15b r 15b r 15: 0 16: 1 16: 1 16: 0 17: 0 18: 1 18: 1 18: 0 19: 0 20: 0

12

MAGEs

Fig. 5. Distribution and concentration of MAGEs in SPOM from ESAO surface waters (Panel A). Panel B illustrates distributions and concentration of MAGEs in SPOM collected from different depths at ESAO stations (St. 3).

Table 2 Abundance (ng/g) of MAGEs in sediments from the ESAO. Depth (cm)

1 3 5 7 9

Total MAGEs abundance St. 0

St. 1

St. 3

St. 8

St. 10

St. 6

St. 13

86 190 370 0 80

240 260 140 86 73

120 72 31 17 15

401 301 306 120 210

510 170 180 103 86

310 150 110 34 67

830 350 320 360 170

water column. This has been observed under oxic conditions in the Cariaco Trench (Hastings and Emerson, 1988), and methane and S2 (microbially produced under anoxic conditions) have been detected in oxygenated waters (Brooks et al., 1981; Burke et al., 1983; Cutter and Krahforst, 1988). Moreover, sulfate reduction might occur in O2 free environments (e.g. oceanic O2 minimum zones) via a cryptic sulfur cycle (Canfield et al., 2010) by which the sulfide produced by sulfate reduction is immediately oxidised to sulfate, resulting in sulfate reduction being overlooked. In addition, SRB have been quantified in oxic regions of the upper water column in a stratified fjord in Denmark (Teske et al., 2002). These

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Fig. 6. Distribution and concentration (ng/g) of MAGEs in surface sediments (upper 10 cm) from the ESAO. A, branched C15 MAGE; b, branched C15 MAGE; c, C15:0 MAGE; d, C16:1 MAGE; e, C16:1 MAGE; f, C16:0 MAGE; g, C17:1 MAGE; h, branched C17 MAGE; I, branched C17 MAGE; j, C17:0 MAGE; k, C18:1 MAGE: l, C18:1 MAGE; m, C18:0 MAGE; n, C19:0 MAGE; o, C20:0 MAGE.

observations have led to the suggestion that highly reducing (anoxic) microenvironments can exist within marine snow or in aggregates that host communities of anaerobic microbes (Fukui and Takii, 1990; Sieburth, 1993; Shanks and Reeder, 1996). The presence of MAGEs in SPOM and sinking particles in SO and ESAO waters could therefore be the result of SRB living in anoxic microenvironments within such settings. Although that possibility cannot be excluded, the widespread occurrence of MAGEs in SPOM from multiple sites and at relatively high concentration suggests that they are associated with aerobic organisms. Their source could be phytoplankton or aerobic bacteria, likely heterotrophic bacteria, given that total MAGE concentration correlated with phytoplankton biomarker concentrations (sum of alkenones and phytol; r 0.66, n = 16, p < 0.05; Fig. 2c and below) and that bacterial growth is controlled by OM production. In addition, a poor correlation between total MAGE concentration and cholesterol concentration in SPOM (r 0.39, n = 14, p < 0.17; Fig. 2a) suggests that the former are not produced by animal sources (e.g. zooplankton). 4.1.2. Sources of MAGEs in sediments Given their occurrence in SPOM and evidence from sediment traps that they are exported from the photic zone, a planktonic source of sedimentary MAGEs is likely. Therefore, and given the lack of evidence for sulfate reduction in underlying sediments (e.g. detection of HS), we expect SPOM and sedimentary MAGE distributions to be similar. However, this was not the case at any of our sites. MAGEs in south Atlantic surface waters ranged from C12 to C20 and the distributions were dominated by either the saturated C14 or C16 MAGE (Fig. 5) but, in underlying surface sediments, only the C15 to C20 components were detected and their distributions were dominated by either the monounsaturated C16

or the saturated C17 component (except St. 0, where the MAGE distribution was still dominated by the saturated C16 component). MAGEs in SPOM from north of the plateau (SO) were dominated by the C14 to C20 components, with the distributions dominated by the monounsaturated C14 or the diunsaturated C16 MAGE (Fig. 7). Sediment trap and surface sediment MAGEs (upper 2 cm) were also dominated by C16, but these were the only MAGEs detected. SPOM from south of the Plateau (SO) contained MAGEs ranging from C15 to C19 and the distribution was dominated by the diunsaturated C16 MAGE. In sediments, however, only C15 and C16 MAGEs were detected and the branched C15 component dominated. Thus, the distributions of MAGEs in surface sediments were significantly different from those in SPOM from overlying water in both the SO and ESAO. This could reflect the fact that the sediments represent inputs derived from hundreds of years, in comparison with the temporally restricted SPOM sampling (e.g. SAPS filtering water for a few h; Hernandez-Sanchez et al., 2012). However, such a sampling bias seems unlikely at all of our sites, especially in the offshore ESAO water, where oceanographic conditions are relatively stable (Hardman-Mountford et al., 2003). Preferential degradation of certain MAGEs could result in different sedimentary distributions; specifically, this would require: degradation of C14:0 and C16:0 MAGEs (C18:0 at St. 0.75) and preferential preservation of C16:1 and C17:0 MAGEs (C16:0 at St. 0.75) in the ESAO; degradation of C14:0 MAGE and preservation of C16:0 and C16:2 MAGEs north of the Crozet Plateau (SO); and degradation of C16:2 MAGEs relative to branched C15:0 MAGEs south of the Plateau (SO). Preferential degradation of MAGEs (relative to each other) has not been previously assessed, but it is difficult to understand why the same compounds would behave differently among our study areas. For example, MAGE

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M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

2.0

Crozet North

Crozet South

20

a

F 15

1.0

10

0.5

5

0 12

0

ng/l

1.5

b

10 8

G

4

ng/l

3

6 2 4 1 2 0

c

14: 15b0 15br 15: r 16:0 2 16: 2 16: 1 16: 1 16: 0 17: 1 17b 17: r 0 18: 18:2 1 18: 1 18: 1 18: 1 18: 1 18: 0 19: 0 20: 0

0

800

ng/l

600

400

Monounsaturated MAGEs 200

Branched MAGEs Saturated MAGEs

0 70

d

60

ng/l

50 40 30 20 10 0 6

e

ng/l

5 4 3 2 1

14: 15b0 15br 15: r 0 16: 2 16: 2 16: 1 16: 1 16: 0 17: 1 17b 17: r 0 18: 18:2 1 18: 1 18: 1 18: 1 18: 0 19: 0 20: 0

0

Fig. 7. Distribution and concentration of MAGEs in SPOM from Southern Ocean surface water (ca. 50 m depth). Crozet North refers to the Fe fertilised area north of the Crozet Plateau, whereas Crozet South corresponds to the HNLC area to the south. Panels A to E represent SPOM samples from north of the Crozet Plateau at station M5 (panels A, D and E), M3 (panel B) and M10 (panel C). Panels F and G represent two SPOM samples collected south of the Plateau at station M6.

distributions were similar in overlying water north and south of the plateau (with C16:2 dominating; Fig. 7), but C16:2 MAGEs seemed to be preferentially degraded south of the Plateau and preserved north of it. O2 penetration depth, and consequently O2 exposure time within sediments as greater south of the Plateau (Homoky et al., 2011). Under these circumstances more intense degradation of the C16:2 MAGE would be expected south of the plateau, consistent with our observations. However, it cannot account for the absence of the C15 MAGE north of the Plateau

and its presence south of the Plateau. In addition, preferential preservation of C16:1 and C16:2 MAGEs seemed to occur at most stations sampled in our ESAO transect, regardless of showing significant differences in exposure times to O2 within sediments (e.g. Stations 1, 3, 4.5 and 6 vs. Stations 0, 0.5 and 0.75; Fig. 8; note that MAGE distributions differed between stations in surface water). Therefore, it seems unlikely that differences in MAGE distributions between SPOM and underlying sediments were solely due to differences in biomarker reactivity.

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Table 3 Abundance (ng/l) of MAGEs in sinking particles and sediments from the Crozet Plateau (SO). MAGE

Sediment traps northa S1

C15 Br C15 Br C15:0 C16:2 C16:1 C16:0 a

S2

Sediments

S3

S4

S5

North

South 27 12 12

495 1.83

0.81

0.0054

13

26 26

12

S1–S5 refer to different sediment trap cups.

The differences could instead (or additionally) arise from a diverse population of aquatic MAGE-producing organisms, all with different compound assemblages as well as different export characteristics. For example, C16:2 MAGE dominated SPOM south of the Plateau, but poor export efficiency of the source organism could explain its absence from sediments. Degradation of sinking OM in the water column (Sherr and Sherr, 1996; Azam, 1998; Rivkin and Legendre, 2001) depends on its residence time, which is a function of sinking rate and, by extension, a wide range of factors, including, for example, aggregation of OM with biogenic or lithogenic material (Ittekkot, 1993; Armstrong et al., 2001). Thus, differential degradation of MAGEs could arise from variable source associations and, given our limited understanding of MAGE occurrence in various phytoplankton and bacteria, this explanation cannot be excluded. To further evaluate a potential planktonic input of sedimentary MAGEs, we compared their concentration with those of

co-occurring phytoplankton biomarkers in the ESAO transect (Fig. 2c–f), including: long chain alkenones (total C37 and C38 components), derived from coccolithophorids (Volkman et al., 1980, 1995; Marlowe et al., 1984); C28 1,14 diol and C29 hydroxymethyl alkanoate, derived from Proboscia spp. diatoms (Sinninghe Damsté et al., 2003); and phytol, derived from all phytoplankton (Fig. 2d-f). In oxygenated sediments, total concentration of MAGEs correlated with phytoplanktonic biomarker concentrations (r 0.77, n = 11, p < 0.05; Fig. 2e); however, there was poor correlation between these two variables in underlying sediments (r 0.44, n = 15, p < 0.15; Fig. 2f). O2 concentration in these intervals was below detection, but sulfide was also absent, such that they are considered to be nitrogenous–ferruginous sediments (dissolved NO 3 and Fe in porewater; Homoky et al., 2013). Thus, it is possible that some of the MAGEs in oxic sediments are at least partially sourced from overlying water and have a planktonic origin. Because of the difference between SPOM and sedimentary MAGE distributions, however, we suggest that there is likely an additional sedimentary source of MAGEs in both the oxic, and especially the nitrogenous– ferruginous sediments. One potential source of MAGES is fauna, which are able to live within oxic nitrogenous–ferruginous sediments (Fig. 8; Homoky et al., 2013). However, total MAGE and cholesterol concentrations were even more poorly correlated in the sediments than they were in water (r 0.35, n = 27, p < 0.1; Fig. 2b), suggesting that fauna were not the main source of MAGEs within the sediments. Instead, we suggest that, consistent with previous reports, sedimentary bacteria are an important additional source of MAGEs (Hinrich et al., 2002; Teske et al., 2002). Consistent with this is the increased proportion of branched components in sediments relative to SPOM. Branched FAs have been widely used as markers

O2concentration (µmol L-1) 100

0

200

100

0

300

200

300

100

0

200

300

0

St 6

Oxic

St 4.5

St 3

Nitrogenous-ferruginous

1

Depth (cm)

2 3 4 5

O2concentration (µmol L-1)

6

100

0

200

300

0

St 0.75

St 1

St 0.5

St 0

Depth (cm)

1 2 3 4 5 6 0

200

400

600

800

0

200

400

600

800

0

200

400

600

800

0

200

400

600

800

MAGEs concentration (ng/g) Fig. 8. Total MAGE abundance for the top 6 cm in sediments from the ESAO. Grey bars represent abundance in sediments with porewater O2 concentration > 0 (labelled as oxic) and black bars indicate abundances in sediments with porewater O2 concentration zero (labelled as ‘‘nitrogenous–ferruginous’’; Homoky et al., 2013) and the black line shows O2 concentration. Only the first 6 cm are shown as O2 measurements were only carried out to that depth.

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13

for bacteria in sediments (Leo and Parker, 1966; Cranwell, 1973, 1974; Kaneda, 1991; Fang and Barcelona, 1998), given that they are needed by many bacteria (e.g. SRB; Boetius et al., 2000; Orphan et al., 2001) to maintain optimal membrane fluidity (Kaneda, 1991). Thus, it is likely that branched MAGEs are also of bacterial origin. A likely bacterial source is SRB, as argued for a variety of other settings (Hinrich et al., 2002; Orphan et al., 2001; Teske et al., 2002). However, the presence of MAGEs in both oxic and nitrogenous–ferruginous sediments (Fig. 8), where there is no evidence of sulfate reduction (Homoky et al., 2013) suggests that SRB are not the sole source of sedimentary MAGEs. Moreover, differences in MAGE distributions between oxic and nitrogenous–ferruginous sediments in the ESAO indicate a complex mixture of sources; for example, proportions of unsaturated C16 and branched components increased with sediment depth and into nitrogenous– ferruginous horizons at some stations (Fig. 6). Thus, we suggest that there are at least three different sources of sedimentary MAGEs different from SRB: those associated with export from the photic zone, those associated with aerobic organisms and those associated with anaerobic organisms associated with nitrogenous–ferruginous environments, (perhaps NO 3 , Mn or Fe reducing microorganisms). 4.2. Distribution of MAGEs under different oceanic regimes: a temperature control? MAGEs were present in SPOM collected across three very different productivity regimes: (i) a high productivity region (40°S) in the ESAO, (ii) a high productivity region in the SO fuelled by natural Fe fertilisation (Seeyave et al., 2007) and (iii) a HNLC area of the SO limited by Fe. The productivity regimes north and south of the plateau (SO) are very different, with a phytoplankton community dominated by diatoms and Phaeocystis spp. (Poulton et al., 2007) and with a microbial activity higher than that south of the Plateau (Zubkov et al., 2007). However, the difference did not manifest itself in a difference in the distribution of MAGEs, which was broadly similar between the two areas (Fig. 7). North of the plateau, MAGEs ranged from C14 to C20 and the distribution was dominated either by the C14:0 or C16:2 components. South of the Plateau, MAGEs ranged from C15 to C19 and the distribution was also dominated by the C16:2 component. Therefore, Fe fertilisation – and the dramatic changes in phytoplankton assemblage that accompany it (Seeyave et al., 2007; Hernandez-Sanchez et al., 2010) – does not seem to strongly affect the distribution of these compounds (although it does affect their concentration). A very different distribution of MAGEs was observed in the ESAO waters. There, MAGEs varied from C12 to C20 and the distribution was dominated by the saturated C14 and C16 components. Thus, the ESAO MAGEs have both slightly longer chain length (e.g. average chain length varying from 15.3 to 16.6 in ESAO SPOM vs. 15.4–16.2 in SO SPOM) and a lower degree of unsaturation. We have calculated the unsaturation index for MAGEs in both oceanic regimes (ESAO and SO) as the concentration of total unsaturated compounds relative to total saturated compounds, and the difference between SO and ESAO SPOM is striking. In the SO, the index ranges from 0.4 to 12 and in the ESAO from 0.08 to 0.4. Given the relatively similar values throughout the Crozet regime, it seems unlikely that the high degree of unsaturation in SO waters vs. ESAO waters reflects the phytoplankton community, and it instead appears to be an adaptation to the environmental conditions. Organisms have the ability to regulate the fluidity of their membranes, depending on environmental conditions, including varying the degree of unsaturation of the constituent FAs or remodelling their polar head groups. Under low temperatures, membrane lipid bilayers change from a fluid state to a non-fluid

11

array of FA chains (de Mendoza and Cronan, 1983; Vigh et al., 1998); in order to maintain their fluidity, membrane components typically have a greater number of double bonds, as unsaturated FAs form a solid state at much lower temperatures (Cronan et al., 1987). This is only true for cis unsaturation, due to the rigid kink of the cis double bond that results in poorer packing of unsaturated FAs (Cronan et al., 1987; Mansilla et al., 2004). Although we did not determine the stereochemistry of MAGEs, a similar increase in the degree of unsaturation with lower temperature has been noted in the lipids of other planktonic organisms, especially ketones synthesized by haptophyte algae (alkenones; e.g. Prahl and Wakeham, 1987; Müller et al., 1998). This has led to their widespread use as a proxy for sea surface temperature (Brassell et al., 1986). Although the physiological function of alkenones is still uncertain, they are believed to play a role in controlling metabolic storage and/or membrane rigidity (Epstein et al., 2001; Sawada and Shiraiwa, 2004). Sea surface temperature is colder around the Crozet Plateau (located within the PFZ) and, at the time of sampling, was 4–12 °C colder (ranging from 4 to 6 °C; Read et al., 2007) than in the ESAO (from 10 to 18 °C; CTD unpublished data). Thus, we propose that the presence of more unsaturated MAGEs in SO waters is likely related to the lower temperature under this oceanic regime. 5. Conclusions A wide variety of MAGEs, including straight chain and branched chain, as well as saturated and unsaturated components, is widespread in marine settings, occurring in SPOM and sediments collected from SO and ESAO waters and in sinking particles collected from deep SO waters. MAGEs occur in both oxic and nitrogenous–ferruginous sediments. The relationship between them and phytoplankton biomarkers in oxic sediments and their presence in SPOM and sinking particles within an oxic water column indicates an aerobic water column source and suggests that not all sedimentary MAGEs derive from SRB. We note, however, that the difference in distribution between SPOM and sedimentary MAGEs (with the latter containing more branched components) does indicate additional sedimentary sources (likely bacteria). Therefore, the application of MAGEs as biomarkers for sulfate reduction might not be adequate, given that these compounds are ubiquitous in a range of oxic-nitrogenous ferrigenous marine settings. The distributions of MAGEs not only differ between water and different sedimentary regimes but also between different oceanographic settings. This does not appear to be due to differences in primary productivity, because distributions are similar in the Fe-fertilised vs. the HNLC Crozet regions. However, distributions do differ between ESAO and SO SPOM, with unsaturated MAGEs dominating in SO waters. We attribute the greater degree of unsaturation as an adaptation to lower temperature in the much colder (4–12 °C) SO waters. Acknowledgements The authors would like to thank the NERC funded UK-GEOTRACES programme and companion PaleoProxy investigation (NE/F019076/1), CROZEX and benthic Crozet programs (NE/ B502844/1; NER/A/S/2003/00576) and the EU BIOTRACS program (MEST-CT-2004-514262). We are particularly thankful to Prof. R. Mills for conceiving funding and for very useful discussions. We also thank the officers, crew and technical support on RV Discovery cruises D286, D300 and D357. In particular we are grateful to F. Challian for collection of GFF filters during the benthic Crozet cruise D300 and M. Badger for software support. R.D.P acknowledges the Royal Society Wolfson Research Merit Award.

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We also thank C. Zhang and an anonymous reviewer for valuable comments on the manuscript.

Associate Editor—J.K. Volkman References Armstrong, R.A., Lee, C., Hedges, J.I., Honjo, S., Wakeham, S.G., 2001. A new mechanistic model for organic carbon fluxes in the ocean based on the quantitative association of POC with ballast minerals. Nature 367, 260–263. Arning, E.T., Birgel, D., Schulz-Vogt, H.N., Holmkvist, L., Jorgensen, B.B., Larson, A., Peckman, J., 2008. Lipid biomarker patterns of phosphogenic sediments from upwelling regions. Geomicrobiology Journal 25, 69–82. Azam, F., 1998. OCEANOGRAPHY: microbial control of oceanic carbon flux: the plot thickens. Science 280, 694–696. Brassell, S.C., Eglinton, G., Marlowe, I.T., Pflaumann, U., Sarnthein, M., 1986. Molecular stratigraphy: a new tool for climatic assessment. Nature 320, 129– 133. Brooks, J.M., Reid, D.F., Bernard, B.B., 1981. Methane in the upper water column of the Northwestern Gulf of Mexico. Journal of Geophysical Research 86, 11029– 11040. Boetius, A., Ravenschlag, K., Schubert, C.J., Rickert, D., Widdel, F., Gieske, A., Amann,  rgensen, B.B., Witte, U., Pfannkuche, O., 2000. A marine microbial R., Jø consortium apparently mediating anaerobic oxidation of methane. Nature 407, 623–626. Burke, R.A., Reid, D.F., Brooks, J.M., Lavoie, D.M., 1983. Upper water column methane geochemistry in the eastern tropical north Pacific. Limnology and Oceanography 14, 454–458. Canfield, E.D., Stewart, F.J., Thamdrup, B., De Brabandere, L., Dalsgaard, T., Delong, E.F., Revsbech, N.P., Ulloa, O., 2010. A cryptic sulfur cycle in oxygen-minimumzone waters off the Chilean Coast. Science 330, 1375–1378. Carter, H.E., Smith, D.B., Jones, D.N., 1958. A new ethanolamine-containing lipid from egg yolk. Journal of Biological Chemistry 232, 681–694. Clarke, A., 1984. Lipid composition of two species of Serolis (Crustacea, Isopoda) from Antarctica. British Antarctic Survey Bulletin 64, 37–53. Cranwell, P.A., 1973. Branched chain and cyclopropenoid acids in a recent sediment. Chemical Geology 11, 307–312. Cranwell, P.A., 1974. Monocarboxylic acids in lake sediments: indicators derived from terrestrial and aquatic biota of paleoenvironmental trophic levels. Chemical Geology 14, 1–14. Cronan, J.E., Gennis, R.B., Maloy, S.R., 1987. Cytoplasmic membrane. In: Neidhardt, F.C., Ingraham, J.L., Low, K.B., Magasanik, B., Schaechter, M., Umbarger, H.E. (Eds.), Escherichia coli and Salmonella: Cellular and Molecular Biology. American Society for Microbiology, Washington, D.C., pp. 31–55. Cutter, G.A., Krahforst, C.F., 1988. Sulfide in surface waters of the western Atlantic Ocean. Geophysical Research Letters 15, 1393–1396. De Mendoza, D., Cronan, J.E., 1983. Thermal regulation of membrane lipid fluidity in bacteria. Trends in Biochemical Sciences 8, 49–52. Egge, H., 1983. Mass spectrometry of ether lipids. In: Mangold, H.K., Paltauf, R. (Eds.), Ether Lipids. Academic Press, London, pp. 16–47. Elvert, M., Hopmans, E.C., Treude, T., Boetius, A., Suess, E., 2005. Spatial variations of methanotrophic consortia at cold methane seeps: implications from a highresolution molecular and isotopic approach. Geobiology 3, 195–209. Epstein, B.L., D’Hondt, S., Hargraves, P.E., 2001. The possible metabolic role of C37 alkenones in Emiliania huxleyi. Organic Geochemistry 32, 867–875. Fang, J., Barcelona, M., 1998. Biogeochemical evidence for microbial community change in a jet fuel hydrocarbons-contaminated aquifer. Organic Geochemistry 29, 899–907. Fischer, W., Hartman, R., Peterkatalinic, J., Egge, H., 1994. (S)-2-Amino-1,3propanediol-3-phosphate-carrying diradylglyceroglycolipids. Novel major membrane lipids of Clostridium innocuum. European Journal of Biochemistry 223, 879–892. Fricke, H., Gercken, G., Oehlenshlager, J., 1986. 1-O-alkylglycerolipids in Antarctic krill (Euphausia superb Dana). Comparative Biochemistry and Physiology B 85, 131–134. Fukui, M., Takii, S., 1990. Survival of sulphate-reducing bacteria in oxic surface sediments of a seawater lake. FEMS Microbial Ecology 73, 317–322. Goossens, H., Rijpstra, W.I.C., Düren, R.R., de Leeuw, J.W., Schench, P.A., 1986. Bacterial contribution to sedimentary organic matter, a comparative study of lipid moieties in bacteria and recent sediments. Organic Geochemistry 10, 683– 696. Hallgren, B., Larsson, S., 1962. The glyceryl ethers in the liver oils of elasmobranch fish. Journal of Lipid Research 3, 31–38. Hallgren, B., Niklasson, A., Stallberg, G., Thorin, H., 1974. On the occurrence of 1-O(2-methoxyalkyl) glycerols and 1-O-phytanylglycerol in marine animals. Acta Chemica Scandinavica B 28, 1035–1040. Hardman-Mountford, N.J., Richardson, A.J., Agenbag, J.J., Hagen, E., Nykjaer, L., Shillington, F.A., Villacastin, C., 2003. Ocean climate of the South East Atlantic observed from satellite data and wind models. Progress in Oceanography 59, 181–221. Hastings, D., Emerson, S., 1988. Sulfate reduction in the presence of low oxygen levels in the water column of the Cariaco Trench. Limnology and Oceanography 33, 391–396.

Hayasaki, K., Kishimura, H., 2002. Amount and composition of diacyl glycerol ethers in various tissue lipids of the deep sea squid Berruteuthis magister. Journal of Oleo Science 51, 523–530. Hernandez-Sanchez, M.T., Venables, H.M., Mills, R.A., Wolff, G., Fisher, E.H., Holtvoeth, J., Leng, M.J., Pancost, R.D., 2010. Productivity variations around the Crozet Plateau: a naturally Fe fertilised area of the Southern Ocean. Organic Geochemistry 41, 767–778. Hernandez-Sanchez, M.T., Holtvoeth, J., Mills, R.A., Wolff, G.A., Pancost, R.D., 2012. Signature of organic matter exported from naturally Fe-fertilised oceanic waters. Deep Sea Research I 65, 59–72. Hinrich, K.U., Summons, R.E., Orphan, V., Sylva, S.P., Hayes, J.M., 2002. Molecular and isotopic analysis of anaerobic methane-oxidizing communities in marine sediments. Organic Geochemistry 31, 1685–1701. Homoky, W.B., Hembury, D.J., Hepburn, L.E., Mills, R.A., Statham, P.J., Fones, G.R., Palmer, M.R., 2011. Iron and manganese diagenesis in deep sea volcanogenic sediments and the origins of pore water colloids. Geochimica et Cosmochimica Acta 75, 5032–5048. Homoky, W.B., John, S.G., Conway, T.M., Mills, R.A., 2013. Distinct iron isotopic signatures and supply from marine sediment dissolution. Nature Communications 4, 2143. Huber, R., Wilharm, T., Huber, D., Trincone, A., Burggraf, S., Rachel, R., Rockingerm, I., Fricke, H., Stetter, K.O., 1992. Aquifex pyrophilus gen. Nov., sp. Nov., represents a novel group of marine hyperthermophylic hydrogen-oxidizing bacteria. Systematic and Applied Microbiology 15, 340–351. Imbs, A., Demina, O., Demidora, D., 2006. Lipid class and fatty acid composition of the boreal soft coral Gersemia rubiformis. Lipids 41, 721–725. Ittekkot, V., 1993. The abiotically driven biological pump in the ocean and shortterm fluctuations in atmospheric CO2 contents. Global and Planetary Change 8, 17–25. Jahnke, L.L., Eder, W., Huber, R., Hope, J.M., Hinrichs, K.U., Hayes, J.M., Des Marais, D.J., Cady, S.L., Summons, R.E., 2001. Signature lipids and stable carbon isotope analyses of octopus spring hyperthermophilic communities compared with those of Aquificales representatives. Applied and Environmental Microbiology 67, 5179–5189. Jayasinghe, C., Gotoh, N., Wada, S., 2003. Variation in lipid classes and fatty acid composition of salmon shark (Lamna ditropis) liver with season and gender. Comparative Biochemistry and Physiology 134, 287–295. Kaneda, T., 1991. Iso- and anteiso-fatty acids in bacteria: biosynthesis, function and taxonomic significance. Microbiological Reviews 55, 288–302. Koga, Y., Morri, H., 2005. Recent advances in structural research on ether lipids from Archaea including comparative and physiological aspects. Bioscience, Biotechnology, Biochemistry 69, 2019–2034. Kamio, Y., Kanegasaki, S., Takahashi, H., 1969. Occurrence of plasmalogens in anaerobic bacteria. Journal of General and Applied Microbiology 15, 439–451. Karnovsky, M.L., Rapson, W.S., Blank, M., 1946. South African fish products. Part XXVII. The occurrence of a-glyceryl ethers in the unsaponifiable fraction of natural fats. Journal of the Society of Chemical Industry 65, 425–428. Kates, M., 1993. Membrane lipids of Archaea. In: Kates, M., Kushner, D.J., Mathesan, A.T. (Eds.), The Biochemistry of Archaea (Archaeobacteria). Elsevier Science Publishers, Amsterdam, pp. 261–295. Langworthy, T.A., Pond, J.L., 1986. Membranes and lipids of thermophiles. In: Brock, T.D. (Ed.), Thermophiles: General, Molecular and Applied Microbiology. Wiley, New York, pp. 107–135. Langworthy, T.A., Holzer, G., Zeikus, J.G., Tornabene, T.G., 1983. Iso- and anteisobranched glycerol diethers of the thermophilic anaerobe Thermodesulfotobacterium commune. Systematic and Applied Microbiology 4, 1–17. Leo, R.G., Parker, P.L., 1966. Branched chain fatty acids in sediments. Science 152, 649–650. Lutjeharms, J.R.E., Meeuwis, J.M., 1984. The extent and variability of South-East Atlantic upwelling. South African Journal of Marine Science 5, 51–62. Magnusson, C.D., Haraldsson, G.G., 2011. Ether lipids. Chemistry and Physics of Lipids 164, 315–340. Mangold, H.K., Weber, N., 1987. Biosynthesis and biotransformation of ether lipids. Lipids 22, 789–799. Mansilla, M.C., Cybulski, L.E., Albanesi, D., de Mendoza, D., 2004. Control of membrane lipid fluidity by molecular thermosensors. Journal of Bacteriology 186, 6681–6688. Marlowe, I.T., Green, J.C., Neal, A.C., Brassell, S.C., Eglinton, G., Course, P.A., 1984. Long chain (n-C37–C39) alkenones in Prymnesiophyceae. Distribution of alkenones and other lipids and their taxonomic significance. European Journal of Phycology 19, 203–216. Müller, P.J., Kirst, G., Ruhland, G., von Storch, I., Rosell-Melé, A., 1998. Calibration of the alkenone paleotemperature index UK0 37 based on core-tops from the eastern South Atlantic and the global ocean (60°N–60°S). Geochimica et Cosmochimica Acta 62, 1757–1772. Nagan, N., Zoeller, R.A., 2001. Plasmalogens: biosynthesis and functions. Progress in Lipid Research 40, 199–229. Ollivier, B., Hatchinkinan, C.E., Prensier, G., Guezennce, J., Garcia, J.L., 1991. Desulfohalobium retbaense gen. nov., sp. Nov., a halophilic sulfate-reducing bacterium from sediments of a hypersaline lake in Senegal. International Journal of Systematic Bacteriology 41, 74–81. Orphan, V.J., Hinrichs, K.U., Ussler, W., Paull, C.K., Taylor, L.T., Sylva, S.P., Hayes, J.M., DeLong, E.F., 2001. Comparative analysis of methane-oxidizing archaea and sulphate-reducing bacteria in anoxic marine sediments. Applied Environmental Microbiology 67, 1922–1934.

M.T. Hernandez-Sanchez et al. / Organic Geochemistry 66 (2014) 1–13 Pancost, R.D., Sinninghe Damsté, J.S., de Lint, S., van der Maarel, M.J.E.C., Gottschal, J.C., the Medinaut Shipboard Scientific Party, 2000. Biomarker evidence for widespread anaerobic methane oxidation in Mediterranean sediments by a consortium of methanogenic Archaea and Bacteria. Applied and Environmental Microbiology 66, 1126–1132. Pancost, R.D., Bouloubassi, I., Aloisi, G., Sinninghe Damsté, J.S., the Medinaut Shipboard Scientific Party, 2001. Three series of non-isoprenoidal dialkyl glycerol diethers in cold-seep carbonate crusts. Organic Geochemistry 32, 695–707. Pancost, R.D., Pressley, S.J., Coleman, J.L., Benning, L.G., Mountain, B.W., 2005. Lipid biomolecules in silica sinters: indicators of microbial biodiversity. Environmental Microbiology 7, 66–77. Pancost, R.D., Pressley, S., Coleman, J.M., Talbot, H.M., Kelly, S.P., Farrimond, P., Schouten, S., Benning, L.G., Mountain, B.W., 2006. Composition and implications of diverse lipids in New Zealand geothermal sinters. Geobiology 4, 71–92. Planquette, H., Statham, P.J., Fones, G.R., Charette, M.A., Moore, C.M., Salter, I., Nédélec, F.H., Taylor, S.L., French, M., Baker, A.R., Mahowald, N., Jickells, T.D., 2007. Dissolved iron in the vicinity of the Crozet Islands, Southern Ocean. Deep Sea Research Part II 54, 1999–2019. Prahl, F.G., Wakeham, S.G., 1987. Calibration of unsaturation patterns in long chain ketone compositions for paleotemperature assessment. Nature 330, 367– 369. Pollard, R.T., Salter, I., Sanders, R.J., Lucas, M.I., Moore, C.M., Mills, R.A., Statham, P.J., Allen, J.T., Baker, A.R., Bakker, D.C.E., Charette, M.A., Fielding, S., Fones, G.R., French, M., Hickman, A.E., Holland, R.J., Huges, J.A., Jickells, T.D., Lampitt, R.S., Morris, P.J., Nedelec, F.H., Nielsdottir, M., Planquette, H., Popova, E.E., Poulton, A.J., Read, J.F., Seeyave, S., Smith, T., Stinchcombe, M., Taylor, S., Thomalla, S., Venables, H.J., Williamson, R., Zubkov, M.V., 2009. Southern Ocean deep-water carbon export enhanced by natural iron fertilization. Nature 457, 577–580. Poulton, A.J., Moore, M., Seeyave, S., Lucas, M.I., Fielding, S., Ward, P., 2007. Phytoplankton community composition around the Crozet Plateau, with emphasis on diatoms and Phaeocystis. Deep Sea Research Part II 54, 2085–2105. Read, J.F., Pollard, R.T., Allen, J.T., 2007. Sub-mesoscale structure and the development of and eddy in the Subantarctic Front north of the Crozet Islands. Deep Sea Research II 54, 1930–1948. Rivkin, R.B., Legendre, L., 2001. Biogenic carbon cycling in the upper ocean: effects of microbial respiration. Science 291, 135–142. Rütters, H., Sass, H., Cypionka, H., Rullkötter, J., 2001. Monoalkylether phospholipids in the sulfate reducing bacteria Desulfosarcina variabilis and Desulforhadus amnigenus. Archives of Microbiology 176, 435–442. Rütters, H., Sass, H.C., Cypionka, H., Rullkötter, J., 2002. Microbial communities in a Wadden Sea sediment core. Clues from analysis of intact glyceride lipids and released fatty acids. Organic Geochemistry 33, 803–816. Sawada, K., Shiraiwa, Y., 2004. Alkenone and alkenoic acid compositions of the membrane fractions of Emiliania huxleyi. Phytochemistry 65, 1299–1307. Seeyave, S., Lucas, M.I., Moore, C.M., Poulton, A.J., 2007. Phytoplankton productivity and annual community structure in the vicinity of the Crozet Plateau during austral summer 2004/2005. Deep Sea Research II 54, 2020–2044.

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Shannon, L.V., 1985. The Benguela ecosystem, Part I. Evolution of the Benguela, physical features and processes. Oceanography and Marine Biology: An Annual Review 23, 105–182. Shanks, A.L., Reeder, M.L., 1996. Reducing microzones and sulphide production in marine snow. Marine Ecology Progress Series 96, 43–47. Shelton, P.A., Hutchings, L., 1982. Transport of anchovy, Engraulis capensis Gilchrist, eggs and early larvae by a frontal jet current. Journal du Conseil Permanent International pour l’Exploration de la Mer 40, 185–198. Shelton, P.A., Hutchings, L., 1990. Ocean stability and anchovy spawning in the southern Benguela current region. Fishery Bulletin 88, 323–338. Sherr, E.B., Sherr, B.F., 1996. Temporal offset in oceanic production and respiration processes implied by seasonal changes in atmospheric oxygen: the role of heterotrophic microbes. Aquatic Microbial Ecology 11, 91–100. Sieburth, J.M., 1993. C1 bacteria in the water column of Chesapeake Bay, USA. I. Distribution of sub-populations of O2-tolerant, obligately anaerobic methylotrophic methanogens that occur in microniches reduced by their bacterial consorts. Marine Ecology Progress Series 95, 67–80. Sinninghe Damsté, J.S., Rampen, S., Rijpstra, I.W., Abbas, B., Muyzer, G., Schouten, S., 2003. A diatomaceous origin for long-chain diols and mid-chain hydroxyl methyl alkanoates widely occurring in quaternary marine sediments: indicators for high nutrient conditions. Geochimica et Cosmochimica Acta 67, 1339–1348. Teske, A., Hinrich, K.U., Edgcomb, V., de Vera Gomez, A., Kysela, D., Sylva, S.P., Sogin, M.L., Jannasch, H.W., 2002. Microbial diversity of hydrothermal sediments in the Guaymas basin: evidence for anaerobic methanotrophic communities. Applied and Environmental Microbiology 68, 1994–2007. Vigh, L., Maresca, B., Harwood, J.L., 1998. Does the membrane’s physical state control the expression of heat shock and other genes? Trends in Biochemical Sciences 23, 369–374. Volkman, J.K., Eglinton, G., Corner, E.D.S., Forsberg, T.E.V., 1980. Long-chain alkenes and alkenones in the marine coccolithiphorid Emiliana huxleyi. Phytochemistry 19, 2619–2622. Volkman, J.K., Barrett, S.M., Blackburn, S.I., Sikes, E.L., 1995. Alkenones in Gephyrocapsa oceanica: implications for studies of paleoclimate. Geochimica et Cosmochimica Acta 59, 513–520. Wetherbee, B.M., Nichols, P.D., 2000. Lipid composition of the liver oil of deep-sea sharks from the Chatham Rise, New Zealand. Comparative Biochemistry and Physiology B 125, 511–521. Zeng, Y.B., Ward, D.M., Brassell, S.C., Eglinton, G., 1992a. Biogeochemistry of hot spring environments 2. Lipid compositions of Yellowstone (Wyoming, USA) cyanobacterial and Chloroflexis mats. Chemical Geology 95, 327–345. Zeng, Y.B., Ward, D.M., Brassell, S.C., Eglinton, G., 1992b. Biogeochemistry of hot spring environments 3. Apolar and polar lipids in the biologically active layers of a cyanobacterial mat. Chemical Geology 95, 347–360. Zhukova, N.V., 2007. Lipid classes and fatty acid composition of the tropical nudibranch mollusks Chromodoris sp. and Phyllidia coelestis. Lipids 42, 1169– 1175. Zubkov, M.V., Holland, R.J., Burkill, P.H., Croudace, I.W., Warwick, P.E., 2007. Microbial abundance, activity and iron uptake in vicinity of the Crozet Isles in November 2004–January 2005. Deep-Sea Research II 54, 2126–2137.