Deposition of talc — kerolite–smectite — smectite at seafloor hydrothermal vent fields: Evidence from mineralogical, geochemical and oxygen isotope studies

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Chemical Geology 247 (2008) 171 – 194 www.elsevier.com/locate/chemgeo

Deposition of talc — kerolite–smectite — smectite at seafloor hydrothermal vent fields: Evidence from mineralogical, geochemical and oxygen isotope studies Vesselin M. Dekov a,⁎, Javier Cuadros b , Wayne C. Shanks c , Randolph A. Koski d a

Department of Geology and Paleontology, University of Sofia, 15 Tzar Osvoboditel Blvd., 1000 Sofia, Bulgaria b Department of Mineralogy, The Natural History Museum, Cromwell Road, London SW7 5BD, UK c U.S. Geological Survey, 973 Denver Federal Center, Denver, CO 80225, USA d U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025, USA Received 16 July 2007; received in revised form 16 October 2007; accepted 21 October 2007 Editor: D. Rickard

Abstract Talc, kerolite–smectite, smectite, chlorite–smectite and chlorite samples from sediments, chimneys and massive sulfides from six seafloor hydrothermal areas have been analyzed for mineralogy, chemistry and oxygen isotopes. Samples are from both peridotite- and basalt-hosted hydrothermal systems, and basaltic systems include sediment-free and sediment-covered sites. Mgphyllosilicates at seafloor hydrothermal sites have previously been described as talc, stevensite or saponite. In contrast, new data show tri-octahedral Mg-phyllosilicates ranging from pure talc and Fe-rich talc, through kerolite-rich kerolite–smectite to smectiterich kerolite–smectite and tri-octahedral smectite. The most common occurrence is mixed-layer kerolite–smectite, which shows an almost complete interstratification series with 5 to 85% smectitic layers. The smectite interstratified with kerolite is mostly trioctahedral. The degree of crystal perfection of the clay sequence decreases generally from talc to kerolite–smectite with lower crystalline perfection as the proportion of smectite layers in kerolite–smectite increases. Our studies do not support any dependence of the precipitated minerals on the type/subtype of hydrothermal system. Oxygen isotope geothermometry demonstrates that talc and kerolite–smectite precipitated in chimneys, massive sulfide mounds, at the sediment surface and in open cracks in the sediment near seafloor are high-temperature (N 250 °C) phases that are most probably the result of focused fluid discharge. The other end-member of this tri-octahedral Mg-phyllosilicate sequence, smectite, is a moderate-temperature (200–250 °C) phase forming deep within the sediment (∼ 0.8 m). Chlorite and chlorite–smectite, which constitute the alteration sediment matrix around the hydrothermal mounds, are lower-temperature (150–200 °C) phases produced by diffuse fluid discharge through the sediment around the hydrothermal conduits. In addition to temperature, other two controls on the precipitation of this sequence are the silica activity and Mg/Al ratio (i.e. the degree of mixing of seawater with hydrothermal fluid). Higher silica activity favors the formation of talc relative to tri-octahedral smectite. Vent structures and sedimentary cover preclude complete mixing of hydrothermal fluid and ambient seawater, resulting in lower Mg/Al ratios in the interior parts of the chimneys and deeper in the sediment which leads to the precipitation of phyllosilicates with lower Mg contents. Talc and kerolite– smectite have very low trace- and rare earth element contents. Some exhibit a negative or flat Eu anomaly, which suggests Eu

⁎ Corresponding author. Tel.: +359 2 9308 276; fax: +359 2 9446 487. E-mail address: [email protected] (V.M. Dekov). 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.10.022

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depletion in the original hydrothermal fluid. Such Eu depletion could be caused by precipitation of anhydrite or barite (sinks for Eu2+) deeper in the system. REE abundances and distribution patterns indicate that chlorite and chlorite–smectite are hydrothermal alteration products of the background turbiditic sediment. © 2007 Elsevier B.V. All rights reserved. Keywords: Hydrothermal activity; Kerolite–smectite; Oxygen isotope Geothermometry; Smectite; Talc

1. Introduction The formation of clay minerals at seafloor hydrothermal fields has received considerable attention, with most investigations dealing with clays of the smectite and chlorite groups (McMurtry and Yeh, 1981; De Carlo et al., 1983; McMurtry et al., 1983; 1993; Percival and Ames, 1993; Zierenberg and Shanks, 1994; Buatier et al., 1995; Lackschewitz et al., 2000, 2006). Mgcontaining smectites (montmorillonite, stevensite, saponite) and Mg-chlorites (clinochlore) have been studied in detail at seafloor hydrothermal settings (McMurtry and Yeh, 1981; McMurtry et al., 1983; Peter and Scott, 1988; Percival and Ames, 1993; Turner et al., 1993; Lackschewitz et al., 2000; and others). Talc, another Mg-phyllosilicate, has been reported in many papers (Siever and Kastner, 1967; Bischoff, 1969; Spiess et al., 1980; Haymon and Kastner, 1981; Styrt et al., 1981; Zierenberg and Shanks, 1983; Janecky and Seyfried, 1984; Koski et al., 1985; Alt et al., 1987; Percival and Ames, 1993; Turner et al., 1993; Buatier et al., 1995; Marchig and Dietrich, 1996; Lackschewitz et al., 2000; Hannington et al., 2001; Lackschewitz et al., 2006; Ludwig et al., 2006), but has been studied in detail in only a few (Lonsdale et al., 1980; Drits et al., 1989; Zierenberg and Shanks, 1994; D'Orazio et al., 2004). There has been no detailed investigation of kerolite (Cole and Shaw, 1983) and mixed-layer kerolite– smectite (Marchig and Dietrich, 1996; Kuhn et al., 2003) at seafloor hydrothermal fields. In a recent study on the clays from the Grimsey Graben, Dekov et al. (submitted for publication) found that previously reported talc (Hannington et al., 2001; Lackschewitz et al., 2006) and kerolite–stevensite (Kuhn et al., 2003) are in fact kerolite, a relatively uncommon, hydrated and highly disordered variety of talc (Brindley et al., 1977). Careful inspection of previously published X-ray diffraction patterns and calculated structural formulae of talc and stevensite from different seafloor hydrothermal fields reveals that many of these samples are in fact kerolite–smectite. Some authors have recognized talc-like specimens with an expanded d-spacing and have called them hydrated

talc or swelling talc (Drits et al., 1989). We have investigated clay minerals from a wide range of seafloor hydrothermal settings in order to constrain the nature of submarine hydrothermal Mg-phyllosilicates and conditions of formation. 2. Geologic settings We have studied a sample set covering two types of seafloor hydrothermal systems: peridotite-hosted and basalt-hosted, both from sediment-free and sedimentcovered ridges (Table 1; Fig. 1). The only sample from a peridotite-hosted hydrothermal system was taken at the ridge-transform intersection between the St. Paul Fracture Zone (FZ) and Intra-Transform Ridge Segment “A” at the Mid-Atlantic Ridge (D'Orazio et al., 2004). The sample was dredged from the summit area of the Outer Corner High, a highly elevated block of oceanic crust with serpentinized peridotites at the base, gabbros and volcanics above and thin or absent sediment cover. The area of dredging is interpreted as a fossil hydrothermal field (D'Orazio et al., 2004). Vent fluids from peridotite-hosted hydrothermal systems are substantially different from basalt-hosted springs in their pH, alkalinity and composition (Kelley et al., 2001; Douville et al., 2002). We also had only one sample from a basalt-hosted sediment-free hydrothermal system, and that sample was recovered at the margin of an inactive nontronite deposit on the southern summit bench of the Red Seamount, a tholeiitic off-axial seamount at 21°N East Pacific Rise (EPR) (Alt et al., 1987). Samples from basalt-hosted sediment-covered hydrothermal systems were collected from 4 hydrothermal fields: Escanaba Trough, Middle Valley, Guaymas Basin and Grimsey Graben (Table 1; Fig. 1). Hydrothermal fluids of these systems differ (pH, alkalinity, composition) from those of sediment-free hydrothermal systems due to the additional interaction of the fluids with the sediment blanket covering the basaltic axial ridge zone (Von Damm et al., 1985b; Butterfield et al., 1994; Campbell et al., 1994). The Escanaba Trough (Southern Gorda Ridge) is a slow spreading ridge covered by a

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thick (∼ 500 m) hemipelagic and turbiditic sediment blanket (Koski et al., 1994; Zierenberg and Shanks, 1994). A variety of sulfide–sulfate hydrothermal deposits (active and inactive) formed on the flanks of sediment hills that rise above the surrounding seafloor. The concentration of deposits is spatially related to volcanic/intrusive structures that penetrate the sediment blanket and 3 areas have been studied in detail: SESCA, MESCA and NESCA (Koski et al., 1994). Locally, sediment at these areas is hydrothermally altered (Zierenberg and Shanks, 1994). Middle Valley (northern Juan de Fuca Ridge), a rift filled with turbiditic sediments, contains 2 areas of hydrothermal activity (Ames et al., 1993; Percival and Ames, 1993). One of them, the Area of Active Venting (AAV), contains over 20 active sulfate–sulfide vents emanating moderate- to high-temperature (184–276 °C) hydrothermal fluids. Some of the surrounding sediment is hydrothermally altered (Goodfellow and Blaise, 1988). Two of our samples have been collected from active anhydrite-rich and sulfide-bearing chimneys from the AAV (Table 1). Guaymas Basin is part of the rift zone in the Gulf of California that contains 2 overlapping troughs, the Northern and Southern troughs. Both troughs are floored by thick turbidite sediment. In the Southern Trough, the 500 m-thick sediment blanket is intruded by a series of basaltic sills and plugs (Koski et al., 1985). A number of active and inactive sulfide–sulfate deposits cluster on the flat rift floor. The sediment around them is hydrothermally altered (Koski et al., 1985). Two of our samples have been dredged from an active 15 m-high hydrothermal mound in the Southern Trough (Table 1). Grimsey Graben is a sediment-filled pull-apart basin within the Tjörnes Fracture Zone, north of Iceland (Hannington et al., 2001). A shallow-water vent field crowns 2 mound structures in the graben and comprises about 20 active and inactive anhydrite mounds and chimneys. Sediment underlying the vent field is locally hydrothermally altered (Dekov et al., submitted for publication). Five of our samples originate from a gravity core taken from an active hydrothermal mound. 3. Materials and methods We studied 29 samples (sediment, chimney and massive sulfide fragments) from the sites described (Table 1), which were reported to contain “talc” according to previous work. The phyllosilicates in intimate relation to talc and those chemically and

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structurally close to it were investigated as well. Selected macroscopically monomineralic subsamples were air-dried and finely ground in an agate mortar. The mineralogy of these samples was studied by positionsensitive detector X-ray diffractometry (Enraf-Nonius Powder Diffraction System 120 X-ray diffractometer with monochromatic Cu Kα radiation, 40 kV, 35 mA, 0.24 × 5 mm divergence slit and no receiving slit) of random powder mounts, from 2 to 120 °2θ. This type of apparatus collects the diffracted radiation in the entire angular range simultaneously, using an arch-shaped detector. The angle between the incident beam and the sample plane was 2–3°. The samples were spun to maximize the analyzed volume and the counting time was 15 min. We used Si, Y2O3 and C22H44O2·Ag for calibration. The samples with substantial (N2%) amounts of detrital components (quartz, feldspars), FeOOH, and calcite were purified through centrifugation (separation of the b2 μm-size fraction), iron oxide removal following Mehra and Jackson (1960), and calcite dissolution at pH = 4 before further analyses. Oriented air-dried and ethylene glycol-solvated mounts were studied by X-ray diffraction (XRD) (Philips X-ray diffractometer PW 1050 with monochromatic Cu Kα radiation, graphite secondary monochromator, 1° divergence slit, 0.1 mm receiving slit, 42 kV, 42 mA): scans from 2 to 40 °2θ, with steps of 0.05 °2θ, at 8 s/step. Randomly oriented specimens were prepared to identify di- and tri-octahedral clay minerals from the 060 diffraction peaks (57–64 °2θ scans, steps of 0.02 °2θ, at 120 s/step). Some of the clay phases were mixed-layer talc-, kerolite- and chlorite–smectite. We estimated the percentage of swelling smectitic layers in these phases by modeling the XRD patterns of the oriented mounts, using the program MLM by Plançon and Drits (2000). This program calculates XRD patterns of 00l diffraction series of mixed-layer clay minerals with several intercalates. The variables used in the calculations are the relative proportions of the two layer types, type of layer-stacking order (R), size of the coherent scattering domain in the c⁎ direction (defined by the statistical range of number of layers, N, and the mean number of layers, Nmean) and the chemical composition. However, we noticed that Fe content in the calculations not always produced the well-known effects on the simulated patterns, and this precluded its use as a variable. Some of the samples contained more than one clay mineral phase. We estimated their relative abundance by adding the calculated patterns of each phase in the appropriate proportion to match the experimental patterns.

174

Table 1 Description of investigated samples Hydrothermal system type

Hydrothermal system subtype

Locality

Latitude

Longitude

Depth (m)

Sampling device

Sample type

Hand sample description

S2227-15

Peridotitehosted

00°37.1′00°38.6′S

25°33.6′25°34.1′W

2400-2300

Dredge

Sediment

White colloform masses.

20°48′N

109°22′W

1700

Submersible

Sediment

40°46.1′– 40°46.7′N

127°31.1′– 127°31.3′W

3260–3170

Dredge

Sediment

Reddish-white colloform. White, porous, light, massive.

-"-"-"-

St. Paul Fracture Zone — Mid-Atlantic Ridge Intersection, Outer Corner High at the Intra-Transform Ridge Segment “A” Red Seamount, 21°N East Pacific Rise Escanaba Trough, Southern Gorda Ridge, SESCA area, Hill 3170 -"-"-"-

1183–9

Basalt-hosted

Sediment-free

L1-86-NC-15D-2A

-"-

Sedimentcovered

L1-86-NC-15D-2-2 L1-86-NC-15D-2-1 L1-86-NC-15D-2-3B

-"-"-"-

-"-"-"-

-"-"-"-

-"-"-"-

-"-"-"-

-"-"-"-

L1-86-NC-15D-2-4

-"-

-"-

-"-

-"-

-"-

-"-

-"-

-"-

L1-86-NC-15D-2-4grey L1-86-NC-15D-2-8 L1-86-NC-15D-3A L1-86-NC-15D-2-3A L1-86-NC-15D-3-5 L1-86-NC-24D-7-1

-"-"-"-"-"-"-

-"-"-"-"-"-"-

-"-"-"-"-"-"-

-"-"-"-"-"40°45.8′– 40°46.2′N

-"-"-"-"-"127°31.0′– 127°30.8′W

-"-"-"-"-"3230–3210

-"-"-"-"-"-"-

-"-"-"-"-"-"-

L1-86-NC-24D-7-2 L1-86-NC-24D-8-1

-"-"-

-"-"-

-"-"-

-"-"-

-"-"-

-"-"-

-"-"-

-"-"-

L2-86-NC-9G-115-122

-"-

-"-

Escanaba Trough, Southern Gorda Ridge, SESCA area

40°45.4′N

127°32.9′W

3240

Gravity core

-"-

-"-"White colloform crust overgrowing … -15D-2-3A Pinkish-white crust overgrowing massive grey matrix. Massive grey matrix. Light grey. Greyish-white. Light grey massive. Light grey. Light grey sediment fragment of sediment breccia cemented by sulfides. -"Light grey, porous with channels. Grey.

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Sample #

-"-

-"-

40°53.8′N

127°31.1′W

3309

-"-

-"-

-"-

41°00.6′– 41°01.1′N

127°30.0′– 127°30.1′W

3320–3300

Dredge

-"-

Light grey.

48°27.58′N

128°42.59′W

2460

Submersible

Chimney

48°27.35′N

128°42.59′W

-"-

-"-

-"-

Greyish-green colloform kerolite overgrowing anhydrite from the inner chimney wall. Greenish-grey colloform kerolite overgrowing 2251-1-1d.

-"-

Escanaba Trough, Southern Gorda Ridge, MESCA area Escanaba Trough, Southern Gorda Ridge, NESCA area Area of Active Venting, Middle Valley, Juan de Fuca Ridge (Heineken Hollow, active chimney) Area of Active Venting, Middle Valley, Juan de Fuca Ridge (Dead Dog Mound, active chimney) -"-

L2-86-NC-16D-3

-"-

-"-

2251-2-3

-"-

-"-

2251-1-1a

-"-

-"-

2251-1-1d

-"-

-"-

-"-

-"-

-"-

-"-

2251-1-1b

-"-

-"-

-"-

-"-

-"-

-"-

-"-

-"-

7D-8b1

-"-

-"-

27°02.7′N

111°22.8′W

2000

Dredge

Massive sulfide

7D-27A1 GCS

-"-

-"-

Southern Trough of Guaymas Basin, Gulf of California -"-

-"-

-"-

-"-

-"-

-"-

SL 347 GC 60-80

-"-

-"-

Grimsey Graben, Tjörnes Fracture Zone (north of Iceland)

66°36.22′N

17°39.32′W

388

Gravity core

Sediment

SL 347 GC SL 347 GC SL 347 GC SL 347 GC

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

-"-"-"-"-

140-160 160-180 200-220 CC

Pale brown lining anhydrite from the inner chimney wall. Pinkish-white lens in the anhydrite chimney substrate. Light grey patch from massive sulphide chunk. Grey colloform surface of dark green mass interstitial to cavernous sulphide aggregate. Pinkish-white colloform lumps forming massive anhydrite-kerolite layers in an active hydrothermal mound. -"-"-"-"-

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

L1-86-NC-27G-77-80

175

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Fig. 1. Location of the sample sites. Legend: A = divergent plate boundaries; B = convergent plate boundaries; C = transform faults; D = sample site at peridotite-hosted hydrothermal system; E = sample site at basalt-hosted sediment-free hydrothermal system; F = sample site at basalt-hosted sedimentcovered hydrothermal system. Sample site numbering: 1 = St. Paul Fracture Zone, Mid-Atlantic Ridge; 2 = Red Seamount, 21°N East Pacific Rise; 3 = Escanaba Trough, Gorda Ridge; 4 = Middle Valley, Juan de Fuca Ridge; 5 = Guaymas Basin, Gulf of California; 6 = Grimsey Graben, Tjörnes Fracture Zone.

Scanning electron microscope (SEM) observations were made on C-coated specimens of fragments of bulk samples and of selected single mineral grains, using a JEOL 5900LV electron microscope with Oxford Instruments INCA EDX detector. The major element composition was determined using a Cameca SX50 electron microprobe (V = 15 keV, I = 13 nA, electron beam diameter of 1 μm). The samples were prepared as polished sections with a completely flat surface. Certified natural oxides and salts were used as standards. Trace elements were analyzed by Inductively Coupled Plasma Mass Spectrometry (ICP-MS) (Varian 800). Total dissolution of the bulk sample was performed by pressurized HF-HClO4-aqua regia attack (Thompson and Walsh, 2003). The accuracy of the analytical results was controlled by measuring the international standard reference materials SCo-1, JSd1, MAG-1, JLK-1. To minimize the isobaric interference of Ba to Eu, the instrument conditions were set to produce a minimal signal of BaO+ while analyzing a standard solution of 10 ppm Ba. While the rate of oxide formation was below 0.5% no further interference correction was applied. Fourier transform infrared

(FTIR) analysis was performed in transmission mode on samples prepared as KBr pellets (sample/KBr ratio of 1/200 mg) using a Perkin Elmer Spectrum One workbench. Spectra were recorded from 4000 to 250 cm− 1, acquiring 8 scans per spectrum with 8 cm− 1 resolution. Oxygen isotope analyses were performed on purified b2 μm separates using a Finnigan Mat 252 mass spectrometer (USGS, Denver). The oxygen isotopic composition of the solids was determined on CO2 prepared from silicates using the BrF5 technique (Clayton and Mayeda, 1963). Data are reported on the per mil (‰) scale relative to the VSMOW standard with analytical precision of approximately ± 0.2‰. Oxygen isotope equilibration temperatures were calculated applying the fractionation equations of Saccocia et al. (in preparation: 1000 ln α talc-water = 11.70 × 10 6 / T 2 − 25.49 × 10 3 /T + 12.48) for talc (also used for kerolite), Savin and Lee (1988) for smectite, and Wenner and Taylor (1973), Zheng (1993), Cole and Ripley (1999) for chlorite. Assuming water–mineral isotope equilibrium, temperatures of formation were calculated using the following δ18O values for the hydrothermal fluid: 0.4‰ for the Escanaba vent fluids

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(Böhlke and Shanks, 1994); − 0.5 to +0.5‰ for the Middle Valley pore fluids (Bent Hill massive sulfide area, ODP drilling; Gieskes et al., 2002); 0.8–1.0‰ for the Guaymas Basin vent fluids (Shanks, 2001); − 0.3‰ for the Grimsey vent fluids (Lackschewitz et al., 2006); 0‰ (seawater value) was chosen for the Red Seamount and St. Paul FZ fluids because there are no data available for them. 4. Results 4.1. Macroscopic description and SEM-EDS Our results indicate the presence of talc, kerolite– smectite, smectite, chlorite–smectite and chlorite. Talc occurs on top of the semi-lithified sediment chunks or in open cracks near the sediment surface. It is white to reddish/pinkish-white, with pearly lustre, very light, porous and colloform (Table 1). Talc overgrows either grey massive sediment forming colloform crusts with varying thickness (Fig. 2a) or green nontronite deposit (Alt et al., 1987). The grey sediment beneath talc is cut by numerous thin fractures perpendicular to the talcsediment contact (Fig. 2a). The sediment adjacent to the fractures has a rusty appearance owing to the presence of FeOOH. Talc forms rosettes and honeycombed aggregates of large, thin, leaf-like crystals (Fig. 3a, b) that are sporadically stained by FeOOH and enclose dispersed sulfide crystals. Pseudomorphs of jarosite and FeOOH after pyrrhotite are often found (Fig. 4d). Kerolite–smectite is white, pinkish-white or grey to greyish-green, colloform, and lines inner walls of anhydrite chimneys (Fig. 2b) and fills interstitial spaces in cavernous massive sulfides (Table 1). It commonly occurs as tiny spherical aggregates (lepispheres) of radiating fine lamellar crystals overgrowing sulfides, atacamite, and sulfates in the chimneys (Fig. 3c, d). After nucleation on sulfide grains and anhydrite crystals, it grows radially resulting in feathery textures and closepacked lepispheres (Fig. 3e). The feathery aggregates of flaky kerolite–smectite crystals form alternating colloform layers separated by sulfidic laminae (Fig. 3f). Prismatic hollow casts of dissolved anhydrite crystals overgrown by kerolite–smectite are often at the kerolite– smectite/anhydrite contact in chimneys (Fig. 3g). Grey massive chloritic sediment occurs in close association with talc as large plate-like crystals of chlorite-like composition (Fig. 4a) and as botryoids of fine lamellar crystals (Fig. 4b) with higher Al content that we identify as chlorite–smectite (see below). Pyrrhotite crystals with fresh edges and faces are frequently scattered in the chloritic sediment (Fig. 4c).

Fig. 2. Photographs of some of the studied samples: (a) talc layer (pinkish-white with silky lustre) covering highly fractured chloritic (grey) sediment; the black film on the talc layer is MnOOH (L1-86NC-15D-2-4); (b) massive colloform kerolite–smectite (grayishgreen) overgrowing massive anhydrite (white) from the inner chimney wall (2251-2-3). Scale bar and alternating blue–white bars = 1 cm.

4.2. XRD XRD analysis of the powder specimens (not shown) allowed us to assess their bulk mineralogy. We then performed XRD analysis of oriented mounts (purified b 2 μm fraction), both air-dried and glycolated, to study the clay minerals in detail (Fig. 5). Where we detected the presence of mixed-layer minerals, we modelled the experimental patterns (only the 00l peak series) to obtain the composition of the mixed-layer phase. Fig. 5 shows the calculated and experimental patterns of some selected samples. The relative intensity of some of the peaks in the modelled patterns did not reproduce well the experimental patterns, as a result of an incorrect response of the program to Fe content in the specimens. In addition, the intensity of the calculated patterns below ∼ 7 °2θ is higher than that of the experimental ones, due to the dependence of the size of the coherent scattering domain on the diffraction angle, which the program cannot simulate (Plançon, 2002). However, the peak positions and width, which determine the composition

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Fig. 4. SEM micrographs (SEI): (a) large plate-like chlorite crystals (L1-86-NC-16D-1); (b) botryoids of fine lamellar chlorite–smectite crystals (L186-NC-16D-1); (c) hexagonal pyrrhotite crystal in chloritic matrix (L1-86-NC-24D-7-1); (d) jarosite pseudomorph on hexagonal pyrrhotite (arrow) in talc matrix (L1-86-NC-16D-1).

of the mixed-layer mineral, are correctly calculated. In a few cases (e.g., Fig. 5d), the pattern of the air-dried specimen was excessively complex due to an inhomogeneous hydration state of the smectite layers and it was not simulated. We therefore relied only on the simulated pattern of the glycolated specimen. When several clay mineral phases were present, their relative abundances were also quantified. The observed patterns indicate a sequence of Mgphyllosilicates ranging from pure talc and Fe-rich talc, through kerolite-rich kerolite–smectite to smectite-rich kerolite–smectite and tri-octahedral smectite (Table 2,

Fig. 5a–f). In many specimens, more than one kerolite– smectite phase was present. The type of layer stacking in kerolite–smectite was always random (R0). The width of the 00l peaks indicates that the size of the coherent scattering domain in the c⁎ direction generally decreases from talc to kerolite-rich kerolite–smectite to smectiterich kerolite–smectite. This value was assessed in the calculated samples using the range of layers in the coherent scattering domain (N) and the mean number of layers (Nmean, Table 2). The large maximum value of N of 300, used in some specimens indicates that they had a very wide range of coherent scattering domains,

Fig. 3. SEM micrographs [secondary electron images (SEI) and back-scattered electron images (BEI)]: (a) rosette of thin talc crystals (SEI; L1-86-NC15D-2-2); (b) large leaf-like talc crystals forming a rosette, note the dark spots of FeOOH on the surface of talc crystals (SEI; L1-86-NC-15D-2-2); (c) lepispheres of fine lamellar mixed-layer kerolite–smectite among massive barite (SEI; 2251-1-1a); (d) close-up of a broken kerolite–smectite lepisphere shown at (c), note the rosette of radiating lamellar crystals and the honeycomb fabric at the surface (SEI); (e) radial growth of flaky kerolite– smectite resulting in a feathery texture after nucleation on the crystal faces of anhydrite (black hollow cast) and on tiny sulfide grains (white spots); note the close packing of kerolite–smectite lepispheres (upper left) (BEI on polished section; 2251-1-1a); (f) colloform layers of radiating kerolite– smectite separated by sulfidic (pyrite, chalcopyrite) laminae (white, spotted); note the radiating kerolite–smectite lepispheres (BEI on polished section; 2251-1-1a); (g) kerolite–smectite (grey) overgrowing anhydrite (white); note the prismatic hollow casts of dissolved anhydrite crystals (BEI on polished section; 2251-1-1a).

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Fig. 5. XRD patterns [experimental (E) and modelled (M)] of selected representative samples: (a) talc (L1-86-NC-15D-2A); (b) Fe-rich talc (1183–9); (c) kerolite-rich kerolite–smectite (7D-8b1); (d) smectite-rich kerolite–smectite (2251-2-3); (e) serpentine and kerolite–smectite (2251-1-1b); (f) smectite (L1-86-NC-27G-77-80); (g) chlorite (L1-86-NC-24D-7-1); (h) chlorite, chlorite-rich chlorite–smectite and kerolite-rich kerolite–smectite (L1-86-NC-15D-2-3A); (i) chlorite–smectite, chlorite and smectite (L1-86-NC-24D-7-2). Indexes 1, 2, and 3 refer to X-ray scans of: oriented airdried mounts, oriented glycolated mounts, and randomly oriented mounts (060 area), respectively. Peak positions in Å. Brt = barite, C = chlorite, H = halite, K–S = kerolite–smectite. Other abbreviations, see Table 2.

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Table 2 Mineralogy of the investigated samples Sample #

Clay minerals (%)

1183–9 L1-86-NC-15D-2A L1-86-NC-15D-2-2 L1-86-NC-15D-2-1

Talc (100) Talc (100) Talc (100) Talc (100) Smectite (traces) Chlorite (traces) Talc (100) Chlorite (traces) Talc (100) Smectite (traces) Chlorite (traces) Kerolite–smectite 1 (66) Kerolite–smectite 2 (34) Kerolite–smectite 1 (62) Kerolite–smectite 2 (38) Serpentine (58) Kerolite–smectite (42) Serpentine (47) Kerolite–smectite 1 (46) Kerolite–smectite 2 (7) Kerolite–smectite (100) Kerolite–smectite (98) Smectite (2) Kerolite–smectite (100) Kerolite–smectite (100) Kerolite–smectite (100) Kerolite–smectite 1 (95) Kerolite–smectite 2 (5) Kerolite–smectite 1 (96) Kerolite–smectite 2 (4) Kerolite–smectite 1 (86) Kerolite–smectite 2 (4) Serpentine (10) Talc (52) Kerolite–smectite (34) Smectite (13) Talc–smectite (27) Chlorite (25) Chlorite–smectite (17) Kerolite–smectite (17) Smectite (14) Smectite (100) Chlorite (traces) Chlorite (100) Talc (?) (traces) Chlorite (38) Chlorite–smectite (31) Kerolite–smectite (31) Chlorite (62) Chlorite–smectite (38) Talc (traces) Chlorite (major) Kerolite (small) Smectite (traces) Chlorite–smectite (87) Kerolite (13)

L1-86-NC-15D-2-3B L1-86-NC-15D-2-4

2251-2-3 2251-1-1a 2251-1-1d 2251-1-1b

7D-8b1 7D-27A1 GCS SL 347 GC 60–80 SL 347 GC 140–160 SL 347 GC 160–180 SL 347 GC 200-220 SL 347 GC CC S2227-15

L1-86-NC-15D-2-8

L1-86-NC-15D-3A

L1-86-NC-27G-77-80 L1-86-NC-24D-7-1 L1-86-NC-15D-2-3A

L1-86-NC-15D-2-4grey

L1-86-NC-24D-8-1

L2-86-NC-9G-115-122

% smectite layers

Ra

80 50 80 50

0 0 0 0

85

0

25 80 5 5

0 0 0 0

5 5 5 10 60 10 50 10 70

0 0 0 0 0 0 0 0 0

80

0

5

0

48 40

1 0

N(Nmean) b

1–20 (8) 1–12 (5) 1–20 (8) 1–12 (5) 1–27 (5) 1–10 (4) 1–12 (5) 1–10 (5) 1–10 (5) 1–300 (10) 1–20 (10) 1–300 (13) 1–300 (13) 1–300 (13) 1–300 (13) 1–10 (5) 1–300 (13) 1–9 (4) 1–20 (10) 1–9 (5) 1–18 (9) 1–300 (16) 1–15 (5) 1–15 (5) 1–20 (15) 1–15 (6) 1–15 (7) 1–20 (15) 1–14 (5)

Type of smectite c

tritritritri-

Comment

Traces of other minerals d

Fe-rich talc End-member talc

FeOOH?

Fe-containing talc

G, Q

Fe-containing talc

G

Little Fe in kerolite Little Fe in kerolite

G G, Crist?, Mrb? G, Mrb, Mc

ditritritritritritritritritritritritritritri-

G, Anh, Q, F, SiO2am? Fe in kerolite Fe in kerolite

Jr, Q Jr, F Py

Fe in kerolite (?) Fe in kerolite (?) Cc Chrysotile (?) Q, F

diditri-(?) didi-(?) ditri(di)-

Crist, F

Fe-rich smectite Mc, Q, G tri-octahedral/ di-,tri-octahedral

5 5

1 0

5

1

1–30 1–15 1–20 1–30 1–15

(10) (6) (15) (10) (6)

Q

ditridiMc, Q

5

1

1–300 (8)

didi-(?) (continued on next page)

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V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

Table 2 (continued) Sample #

Clay minerals (%)

% smectite layers

Ra

N(Nmean) b

Type of smectite c

L2-86-NC-16D-3

Chlorite–smectite (62) Chlorite (38) Chlorite–smectite (46) Chlorite (36) Smectite (18) Chlorite–smectite 2 (41) Chlorite–smectite 1 (27) Chlorite (18) Smectite (14)

10

1

di-

Mc, Q

48

1

di-

G, Q, Mc

15 48

1 1

1–20 (8) 1–200 (15) 1–15 (6) 1–30 (10) 1–15 (6) 1–15 (6) 1–15 (7) 1–200 (15) 1–15 (5)

L1-86-NC-15D-3-5

L1-86-NC-24D-7-2

a b c d

didi-(?) di-(?)

Comment

Expandable chlorite Expandable chlorite

Traces of other minerals d

F, Mc

di-(?)

Layer stacking order. Range (mean) number of layers in the coherent scattering domains. di- = di-octahedral, tri- = tri-octahedral smectite. Anh = anhydrite, Cc = calcite, Crist = cristobalite, F = feldspar, G = gypsum, Jr = jarosite, Mc = mica, Mrb = mirabilite, Py = pyrite, Q = quartz.

reaching large size values. These specimens are talc or kerolite-rich kerolite–smectite. The relative intensity of the 00l peaks permitted assessment of Fe presence in the clay structure, as increasing Fe produces a relative increase of the 001 peak over the 002 and 003 peaks. The most Fe-rich specimen is the talc from Red Seamount (1183–9, Table 2, Fig. 5a). Several samples contain serpentine (Table 2). The 00l peaks of the kerolite–smectite from these samples are rather wide and their intensity low, indicating small coherent scattering domains. The extreme cases show very broad and low-intensity peaks (Fig. 5e). The XRD scan at 57–64 °2θ shows the 060 and adjacent peaks (Fig. 5). The position of the 060 peaks indicates that the smectite interstratified with kerolite is tri-octahedral with only 2 exceptions: 2251-1-1d, where there seems to be a di-octahedral smectite component, and L1-86-NC15D-2-8, where all the smectite appears to be dioctahedral, as there are 2 broad peaks at 1.490 and 1.477 Å (corresponding to a di-octahedral phase or phases) of similar intensity as the 060 peak at 1.530 Å (corresponding to talc and kerolite). Given the mineralogical composition of this specimen (Table 2), all the smectite should be di-octahedral to account for the relative intensities of the mentioned peaks. The width and degree of resolution of the 060 and adjacent hkl peaks indicates the degree of crystal perfection of the specimens, which decreased generally from talc to kerolite–smectite, with lower crystalline perfection as the proportion of smectite layers in kerolite–smectite increased. Sample L1-86-NC-27G-77-80 is a tri-octahedral smectite with no mixed-layering (Table 2, Fig. 5f). This sample is therefore chemically and structurally closer to the talc – kerolite–smectite series but, from the point of view of its occurrence, it is closer to the

chlorite – chlorite–smectite series of the sediment substrate. However, part of the smectite is di-octahedral, as indicated by the broad, weak peak at ∼ 1.50 Å (Fig. 5f3). The samples of the sediment beneath the talc are dominated by chlorite and chlorite–smectite, with a range of composition between pure chlorite and ∼ 50% smectite layers in chlorite–smectite (Table 2, Fig. 5g–i). For these samples, the layer stacking is always ordered of the type R1. For this layer stacking order, there is a super-structure peak at ∼ 34 Å when the relative proportion of chlorite and smectite layers approach 50% (Fig. 5i). One of the specimens showed a d-spacing for the chlorite layers (in chlorite–smectite) of 14.5 Å, higher than the usual 14.15 Å. We refer to this fact in Table 2 indicating that the chlorite component was “expandable”. The coherent scattering domain size in the c⁎ direction (N, Nmean) decreases generally from chlorite to chlorite–smectite. The 060 and adjacent peaks show reflections corresponding to a tri-octahedral chlorite, possibly clinochlore. In this case it was impossible to estimate the di-or tri-octahedral nature of the interstratified smectite using the 060 angular region because chlorite has peaks at 1.53 and 1.50 Å, overlapping the possible positions of the smectite peaks. However, the simulations of the chlorite–smectite patterns allowed a good fit with the experimental data only when di-octahedral smectite was used. The dioctahedral nature of the smectite component was confirmed by the chemical analysis (see below). The relative intensity of the several 00l peaks indicated that the Fe content of these samples, and possibly the distribution of Fe between the octahedral and brucitelike sheets as well, varied from sample to sample. Fig. 5g shows a specimen in which the Fe content was one of the lowest (near-even intensity of the peaks between 5 and 25 °2θ) and Fig. 5h the case in which Fe

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

content was highest (greatest intensity differences for the peaks within 5–25 °2θ). 4.3. Infrared The infrared spectra confirmed the presence and approximate relative abundance of the minerals detected in the XRD patterns. Fig. 6 shows examples of some of

183

them. The spectra also provide an insight into the Fe content of the talc and kerolite–smectite samples, since some show 2-3 OH stretching bands (region 3680– 3640 cm− 1) arising from the various cation combinations: Mg3OH, Mg2FeOH and MgFe2OH (Fig. 6a, b) (Russell and Fraser, 1994). This confirms that the high Fe content measured by chemical methods in some of these talc and kerolite–smectite samples is not an

Fig. 6. IR spectra of some selected samples. Only the spectral regions providing the characteristic features of phyllosilicates are shown. The wide bands in the high wavenumber region that are not labelled correspond to adsorbed water. (a) Fe-rich talc (1183–9); (b) talc (L1-86-NC-15D-2-3B); (c) kerolite-rich kerolite–smectite, serpentine and traces of calcite (S2227-15); (d) chlorite (L1-86-NC-24D-7-1).

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V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

artifact due to FeOOH coating. Furthermore, there is no IR evidence of the presence of FeOOH. Some of the features of the spectrum in Fig. 6c suggest that the serpentine mineral is chrysotile (bands at 3696, 3652, 597 and 309 cm− 1; Russell and Fraser, 1994). 4.4. Geochemistry The major-element composition of the studied samples (Table 3) is in good agreement with the XRD and IR data. End-member talc samples (Table 2) have very low concentrations of Al, Fe, Mn and interlayer cations (Ca, Na, K) (Table 3), except for talc from the Red Seamount (1183–9), which is Fe-rich and contains higher Mn, Zn and Cu (Table 3). The two talc samples that display an Mg2FeOH IR band contain significant Fe concentration (Tables 2, 3). The higher Al content in the kerolite–smectites compared to that in talc is obviously due to the interstratified smectitic component; Al content increases with increasing smectite content (Tables 2, 3). Kerolite–smectites generally contain

enhanced Mn, Zn, Cu and interlayer cation concentrations compared to those of talc. The pure tri-octahedral smectite is slightly depleted in Si and enriched in Al (Table 3). The chlorite composition is within the range of clinochlore. Concentrations of almost all trace and rare-earth elements (REE) (ICP-MS analysis) in talc and kerolite– smectite are lower than those in chlorite, chlorite– smectite and smectite. Only Bi and Mo are higher in the former group (Table 4). Ga, Sr, Cd and Co are comparable in both types of samples. The very high content (0.1– 3.3%) of As, Ba, Cu, Pb and Zn in some of the samples (Table 4), much higher than that detected from clay grains in the microprobe analyses, indicates mineral impurities (sulfides, sulfates, halides) in these bulk samples. The kerolite–smectites have the lowest concentrations of REEs (ΣREE ≈ 1 ppm) (Table 4). Their REE distribution patterns (Fig. 7a) are, generally, characterized by a positive Eu anomaly, negative or no Ce anomaly and slight enrichment in the LREE relative to HREE. The Middle Valley samples and one from

Table 3 Average major-element composition (in wt. %) of the studied samples, from microprobe analysis Sample #

Mineral

1183-9 L1-86-NC-15D-2A L1-86-NC-15D-2-2 L1-86-NC-15D-2-1 L1-86-NC-15D-2-3B L1-86-NC-15D-2-4 2251-2-3 2251-1-1a 2251-1-1d 2251-1-1b 7D-8b1 7D-27A1 GCS SL 347 GC 60-80 SL 347 GC 140-160 SL 347 GC 160-180 SL 347 GC 200-220 SL 347 GC CC S2227-15 L1-86-NC-15D-2-8 L1-86-NC-15D-3A L1-86-NC-27G-77-80 L1-86-NC-24D-7-1 L1-86-NC-15D-2-3A

Talc Talc Talc Talc Talc Talc Kerolite–smectite Kerolite–smectite Serpentine, kerolite–smectite Serpentine, kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Kerolite–smectite Talc, kerolite–smectite Chlorite, chlorite–smectite Smectite Chlorite Kerolite–smectite Chlorite, chlorite–smectite L1-86-NC-15D-2-4grey Chlorite, chlorite–smectite L1-86-NC-24D-8-1 Chlorite L2-86-NC-9G-115-122 Chlorite–smectite L2-86-NC-16D-3 Chlorite–smectite, chlorite L1-86-NC-15D-3-5 Chlorite–smectite, chlorite L1-86-NC-24D-7-2 Chlorite–smectite

a

Total Fe expressed as FeO.

SiO2

Al2O3 MgO

FeOa

MnO TiO2 Cr2O3 ZnO CuO CaO Na2O K2O

61.50 66.68 66.31 66.25 66.21 67.05 60.88 58.57 53.72 57.70 62.53 57.40 65.26 66.41 66.75 66.58 66.32 66.00 56.09 54.56 51.36 37.60 57.38 44.52 37.96 54.87 38.67 51.60 47.87 49.67

0.69 0.26 0.51 0.39 0.28 0.29 3.24 3.64 3.74 1.88 0.95 0.98 1.74 0.54 0.61 0.60 1.54 0.64 13.93 24.14 11.98 27.39 7.90 19.31 23.92 25.15 21.24 21.39 35.58 23.32

15.11 0.66 1.34 0.86 3.34 1.77 6.06 5.81 6.03 2.32 9.95 14.42 0.37 0.35 0.13 0.17 0.29 0.02 3.00 0.76 9.79 4.34 1.25 1.86 7.52 3.66 10.68 10.23 3.01 2.58

0.12 0.03 0.01 0.01 0.04 0.01 0.18 0.16 0.12 0.06 0.42 0.13 0.02 0.03 0.03 0.00 0.04 0.03 0.03 0.01 0.09 0.07 0.02 0.02 0.04 0.09 0.09 0.09 0.16 0.04

21.64 32.04 31.49 32.12 29.85 30.54 28.20 28.76 31.66 36.17 25.21 26.36 32.30 32.43 32.23 32.42 31.48 32.81 25.16 17.22 23.79 28.85 32.91 32.18 29.02 10.72 27.82 12.47 11.54 21.66

0.01 0.01 0.01 0.03 0.05 0.00 0.00 0.15 0.04 0.09 0.03 0.01 0.01 0.03 0.01 0.01 0.01 0.01 0.05 0.69 0.41 0.63 0.09 1.76 1.02 0.58 1.16 0.99 0.92 0.82

0.02 0.03 0.01 0.03 0.00 0.00 0.02 0.02 0.02 0.01 0.01 0.00 0.03 0.02 0.02 0.03 0.01 0.01 0.04 0.05 0.03 0.05 0.03 0.00 0.04 0.05 0.03 0.03 0.02 0.04

0.17 0.08 0.04 0.10 0.06 0.11 0.07 0.04 0.18 0.15 0.24 0.22 0.01 0.00 0.01 0.01 0.01 0.01 0.06 0.07 0.10 0.07 0.07 0.07 0.02 0.06 0.11 0.06 0.05 0.21

0.37 0.03 0.02 0.03 0.00 0.01 0.62 1.70 2.36 0.84 0.13 0.08 0.01 0.00 0.01 0.00 0.00 0.04 0.07 0.05 0.02 0.06 0.02 0.01 0.07 0.03 0.02 0.02 0.03 0.01

0.02 0.02 0.02 0.01 0.03 0.03 0.10 0.16 1.22 0.42 0.05 0.04 0.09 0.06 0.05 0.05 0.09 0.13 0.08 0.87 0.22 0.08 0.02 0.02 0.02 0.04 0.09 0.07 0.03 0.05

0.34 0.11 0.19 0.12 0.10 0.14 0.48 0.81 0.73 0.22 0.45 0.30 0.16 0.11 0.13 0.10 0.17 0.24 1.06 1.16 1.88 0.66 0.26 0.16 0.27 2.15 0.06 0.71 0.56 0.94

0.01 0.04 0.05 0.04 0.04 0.04 0.14 0.17 0.19 0.14 0.02 0.06 0.01 0.02 0.04 0.03 0.04 0.06 0.43 0.41 0.34 0.21 0.05 0.08 0.10 2.61 0.02 2.34 0.24 0.67

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

Guaymas Basin exhibit a weak negative Eu anomaly (Fig. 7a). Samples containing only talc and talc + kerolite–smectite samples have medium REE contents (generally 1 b ΣREE b 100 ppm) (Table 4). They all show similar REE distribution patterns with negative Eu anomalies and LREE enrichment (Fig. 7b). The Red Seamount sample (1183–9) is an exception with very low ΣREE and positive Ce anomaly. Chlorites and chlorite–smectites are the richest in REEs (ΣREE N 100 ppm; Table 4). There is a broad similarity in the shape of chlorite REE patterns (LREE enrichment, flat HREE distributions, negative Eu anomaly), although there are variations in the REE concentrations among the samples (Fig. 7c). The REE abundances in smectite are less than those in chlorite and are approximately in the top range of values of talc (Table 4). The REE distribution pattern for smectite shows LREE enrichment, a flat HREE distribution, and a weak positive Eu anomaly (Fig. 7c). 4.5. Oxygen isotope analysis The δ18O values of the studied clays range from 3.0 to 4.7‰ for talc, from − 0.4 to 7.1‰ for kerolite– smectite, from 3.7 to 4.3‰ for chlorite and chlorite– smectite, and 5.4‰ for smectite (Table 5). The formation temperatures were calculated using the fluid O isotope compositions indicated in the methods section. They vary from 287 to 323 °C for talc, from 226 to 382 °C for kerolite–smectite, from 153 to 207 °C for chlorite and chlorite–smectite, and 257 °C for the one smectite sample (Table 5). The temperature range for chlorite and chlorite–smectite formation was calculated using the equations of Wenner and Taylor (1973) and Cole and Ripley (1999). The equation of Zheng (1993) gave intermediate values. For comparison we have provided δ18O values from the literature (Table 5) for the same minerals from the same submarine hydrothermal fields and have calculated the respective formation temperatures using the fluid O isotope compositions from the methods section. After the inspection of published XRD pattern and structural formula of sample GHF1 from the Grimsey Graben (Lackschewitz et al., 2006) we identify this sample not as saponite, but as kerolite–smectite. We have therefore used the measured δ18O value of this sample as comparative value for kerolite–smectite. 5. Discussion Calculated equilibrium models for mixing between high-temperature hydrothermal solutions and seawater

185

have predicted abundant talc in seafloor hydrothermal deposits (Janecky and Seyfried, 1984). However, talc is scarce in seafloor deposits and this has been attributed to kinetic effects (Janecky and Seyfried, 1984), which might play an important role in constraining the formation of talc during the mixing of hydrothermal fluids with ambient seawater. We find that only a few of all the samples described previously as “talc” are in fact talc. These include 1 sample from the Red Seamount and 5 samples from the Escanaba Trough (Table 2). All other “talc” samples are actually mixed-layer kerolite–smectite showing an almost complete series of interstratification with 5 to 85% smectitic layers. With two exceptions, all the smectite interstratifying kerolite was found to be trioctahedral (Table 2; Fig. 5). We have therefore documented a continuous sequence of tri-octahedral Mg phyllosilicates from talc through kerolite-rich kerolite–smectite to smectite-rich kerolite–smectite with a logical tri-octahedral smectite end-member. Trioctahedral smectite has been reported at several seafloor hydrothermal fields, but assessment of the published XRD patterns and chemical data from Red Seamount (Alt et al., 1987), Escanaba Trough (Zierenberg and Shanks, 1994), Middle Valley (Percival and Ames, 1993), and Grimsey Graben (Lackschewitz et al., 2006) show that some of them are in fact kerolite–smectite and that others display chemical compositions incompatible with tri-octahedral smectite. Mg phyllosilicates in submarine hydrothermal environments form as a result of mixing of hot silica-rich (and Mg-depleted) hydrothermal fluids with cold Mg-rich seawater (Zierenberg and Shanks, 1994; Buatier et al., 1995; Lackschewitz et al., 2000). Higher silica activity favors the formation of talc relative to tri-octahedral smectite. The broad temperature-compositional stability range of talc (Janecky and Seyfried, 1984) is consistent with the observation that 2:1 Mg-phyllosilicates are more common than 1:1 Mg-phyllosilicates in all seafloor hydrothermal fields. Talc precipitation from the hydrothermal fluid would cause proton generation: 3Mg2þ þ 4SiðOHÞ4 ↔Mg3 Si4 O10 ðOHÞ2 þ 4H2 O þ 6Hþ The resultant decrease in pH would lead to alteration of some of the primary sulfide minerals. Hexagonal primary pyrrhotite scattered in the chloritic sediment (Fig. 4c) would convert into jarosite (Fig. 4d) or FeOOH after talc precipitation. Previous investigations (Alt et al., 1987; Drits et al., 1989; Zierenberg and Shanks, 1994) have also documented similar pseudomorphs of jarosite, FeOOH and marcasite after pyrrhotite in talc matrix.

Table 4 Trace and REEs in the investigated samples, from ICP-MS L1-86NC15D2A

L186NC-15D2-2

L186NC15D2-1

L186NC15D2-3B

L186NC15D2-4

22512-3

22511-1a

7D-8b1

7D27A1 GCS

SL 347 GC 60-80

SL 347 GC 140160

SL 347 GC 160180

SL 347 GC 200220

SL 347 GC CC

S222715

L186NC15D2-8

L186NC15D3A

L186NC27G77-80

L186NC24D7-1

L186NC15D2-3A

L186NC15D2-4 grey

L186NC24D8-1

L286NC9G115122

L286NC16D3

L186NC15D3-5

L186NC24D7-2

Mineral a:

T

T

T

T (S,C)

T(C)

T (S,C)

K-S

K-S

K-S

K-S,S

K-S

K-S

K-S

K-S

K-S

K-S, Serp

T,KS,S

T-S,C, C-S, K-S,S

S,(C)

C

C,C-S, K-S

C, C-S (T)

C,K,S

C-S,K

C-S,C

C-S, C,S

C-S, C,S

As (ppm) Ba Be Bi Cd Co Cr Cs Cu Ga Hf Li Mo Nb Ni Pb Rb Sb Sc Sn Sr Ta Th Tl U V W Y Zn Zr La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu ΣREE (Ce/Ce⁎) b (Eu/Eu⁎) c LaN/LuN

28.3 21.2 0.03 16.9 0.09 133 9.25 0.04 16783 21.9 0.01 4.84 70.1 0.09 7.2 4.99 0.85 0.44 0.33 0.61 33.9 0.01 0.01 0.30 0.54 39.6 0.03 0.83 893 2.51 0.04 0.10 0.01 0.03 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00 0.01 0.00 0.24 1.35 0.60 2.71

29.3 9.8 0.09 2.45 0.08 4.29 4.90 0.01 50.5 15.8 0.03 1.50 2.52 0.14 17.9 9.22 0.78 2.32 1.35 2.13 21.5 0.03 0.21 0.01 0.28 10.2 0.15 2.59 17.9 1.02 2.49 6.44 1.00 4.09 0.92 0.20 0.68 0.09 0.46 0.08 0.21 0.03 0.14 0.02 16.9 1.00 0.75 13.6

35.1 21.2 0.14 1.04 0.09 5.21 3.97 0.02 51.4 13.8 0.02 1.58 1.95 0.08 11.4 12.1 1.06 4.41 1.05 2.07 29.6 0.01 0.19 0.01 0.30 6.45 0.12 1.64 19.6 0.78 0.99 2.74 0.41 1.85 0.49 0.08 0.40 0.06 0.33 0.06 0.16 0.02 0.12 0.02 7.73 1.05 0.51 6.17

44.4 47.4 0.15 5.71 0.07 3.31 13.5 0.06 48.8 14.7 0.10 2.51 0.97 0.65 16.3 9.24 1.40 5.79 2.69 4.17 21.6 0.04 0.42 0.03 0.32 12.9 0.27 3.64 14.7 3.68 5.38 14.0 2.17 9.06 2.12 0.77 1.53 0.21 0.91 0.15 0.38 0.04 0.27 0.04 37.0 1.00 1.24 16.1

205 138 0.47 1.16 0.29 15.4 23.0 0.13 167 26.5 0.49 7.92 11.7 1.75 37.2 72.3 3.08 10.3 3.29 6.51 89.9 0.10 1.01 0.27 1.42 116 0.71 14.9 65.2 27.6 11.1 20.3 3.56 15.2 3.84 1.16 3.66 0.58 3.17 0.56 1.60 0.21 1.31 0.19 66.4 0.79 0.93 6.33

194 21.1 0.41 12.5 0.11 8.23 11.7 0.02 95.9 28.9 0.08 8.56 19.9 0.36 24.5 14.2 1.07 9.28 1.62 4.40 31.2 0.02 0.14 0.10 1.48 80.8 0.22 8.81 64.1 7.17 5.29 9.81 1.38 5.52 1.41 0.27 1.40 0.23 1.30 0.23 0.61 0.08 0.45 0.06 28.0 0.87 0.58 8.96

8.32 75.0 0.33 4.81 0.38 5.62 6.06 0.30 7132 7.88 0.02 18.6 4.55 0.10 12.6 43.9 4.78 0.86 0.41 5.08 72.8 0.04 0.04 0.17 0.42 4.44 0.17 0.29 216 0.68 0.13 0.39 0.03 0.12 0.03 0.01 0.04 0.01 0.03 0.01 0.02 0.01 0.03 0.01 0.86 1.44 0.79 3.11

16.7 118 0.26 4.74 0.61 7.78 4.32 0.20 19306 6.19 0.02 12.2 1.89 0.09 16.7 23.7 5.65 2.07 0.27 5.47 128 0.03 0.06 0.18 0.54 3.92 0.17 0.45 266 0.69 0.21 0.45 0.05 0.20 0.05 0.01 0.05 0.01 0.04 0.01 0.03 0.01 0.03 0.01 1.15 1.03 0.87 4.02

0.96 289 0.17 1.65 8.74 8.92 7.38 0.07 2474 3.46 0.10 4.34 7.83 0.24 14.2 29.1 2.46 0.55 0.37 2.38 11.8 0.02 0.16 0.15 43.7 16.5 0.09 0.89 2230 3.22 0.65 1.32 0.17 0.72 0.17 0.14 0.17 0.02 0.11 0.02 0.05 0.01 0.04 0.01 3.59 0.95 2.52 11.3

2.82 55.3 0.03 3.91 31.8 106 1.84 0.04 33177 5.86 0.01 7.72 0.36 0.04 2.0 9.25 1.51 0.22 b 0.007 0.74 4.50 0.01 0.02 0.08 0.13 7.33 0.01 0.30 11505 0.33 0.13 0.25 0.03 0.13 0.03 0.01 0.03 0.01 0.02 0.01 0.01 0.00 0.01 0.00 0.65 0.93 0.63 17.9

0.02 38.1 0.04 0.07 b 0.004 0.70 21.9 0.07 31.6 14.1 0.05 36.4 0.25 0.06 6.34 4.88 0.40 0.07 0.61 1.26 5.42 b 0.01 0.01 0.01 0.68 77.9 0.19 0.09 41.9 b 20 0.10 0.14 0.02 0.07 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.41 0.77 4.71 6.02

0.08 22.4 0.03 0.05 b0.004 0.58 22.0 0.08 9.61 11.9 0.04 17.8 0.15 0.08 8.51 2.12 0.28 0.09 0.57 0.61 4.26 0.04 0.01 0.03 4.29 80.1 0.22 0.06 24.1 b20 0.78 0.04 0.01 0.02 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.90 0.04 6.84 46.4

0.02 69.7 0.05 0.09 b 0.004 1.66 22.9 0.54 26.5 12.3 0.07 27.3 0.12 0.15 11.1 5.13 0.44 0.14 0.72 1.01 6.34 b 0.01 0.02 0.01 2.77 68.8 0.25 0.12 35.7 b 20 0.11 0.18 0.03 0.11 0.03 0.03 0.02 0.01 0.02 0.01 0.01 0.01 0.02 0.01 0.57 0.76 4.10 4.22

1.85 55.4 0.10 0.07 b 0.004 1.19 21.3 0.06 16.0 16.6 0.04 13.0 0.24 0.03 26.7 4.28 0.36 0.16 0.38 0.94 12.0 0.02 0.01 0.01 0.26 63.2 0.17 0.09 42.6 b 20 0.20 0.06 0.01 0.03 0.01 0.03 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.39 0.22 8.21 9.59

0.49 1.36 0.03 0.11 0.01 1.34 20.7 0.24 16.5 6.99 0.05 6.09 0.23 0.06 19.6 6.39 0.23 0.12 0.37 0.86 7.21 b 0.01 0.01 0.01 0.48 91.5 0.18 0.21 28.4 b 20 0.45 0.12 0.02 0.07 0.02 0.01 0.02 0.01 0.03 0.01 0.02 0.01 0.02 0.01 0.78 0.19 1.79 20.0

1.39 3.58 0.28 0.12 0.07 3.16 16.1 0.15 282 8.77 0.13 104 0.82 0.41 39.7 0.27 2.17 0.11 1.57 1.13 55.2 0.02 0.20 0.16 8.21 8.63 0.34 1.96 46.3 b 20 1.57 1.84 0.36 1.39 0.26 0.12 0.28 0.04 0.26 0.06 0.17 0.02 0.16 0.03 6.56 0.58 1.29 6.34

27.0 126 0.68 4.18 0.18 5.69 119 0.08 96.9 9.15 1.62 1.31 0.61 13.4 20.6 21.6 4.64 1.88 14.6 9.00 212 0.93 4.77 0.07 0.64 99.5 0.60 10.7 28.2 48.6 10.6 24.1 3.09 11.7 2.32 0.69 2.12 0.30 1.69 0.31 0.89 0.11 0.68 0.10 58.7 1.02 0.94 11.6

0.90 52.3 0.81 0.70 0.05 1.69 126 0.30 76.2 7.83 2.13 4.43 0.36 15.1 46.2 3.51 5.18 0.31 21.5 5.74 20.0 1.07 10.5 0.04 0.87 191 1.84 17.0 20.3 55.6 23.2 49.4 5.80 22.2 4.36 0.88 3.47 0.45 2.08 0.39 1.18 0.15 1.01 0.15 115 1.02 0.67 16.2

93.2 180 1.00 2.51 0.44 10.7 52.2 1.88 243 15.7 0.98 17.8 8.97 7.42 32.7 7.14 30.1 4.60 14.0 4.51 40.4 0.27 2.56 0.11 6.51 111 0.47 11.7 154 51.6 9.50 19.2 2.35 9.05 1.88 0.62 1.68 0.23 1.24 0.24 0.74 0.10 0.65 0.10 47.6 0.97 1.04 10.3

243 1221 1.08 0.89 0.59 14.6 145 0.14 142 8.37 2.03 19.4 2.33 24.8 73.7 362 2.48 15.9 29.5 19.3 86.3 1.64 13.8 0.25 3.02 200 2.51 22.2 349 66.7 40.9 77.2 8.90 32.6 6.06 1.49 5.31 0.72 3.55 0.65 1.83 0.24 1.48 0.22 181 0.95 0.78 20.1

9.75 3.36 0.88 0.12 0.05 6.57 111 0.06 20.5 7.38 2.73 11.9 1.69 12.8 46.0 2.55 1.17 0.97 24.5 11.2 6.56 0.96 5.31 0.03 0.79 156 2.61 3.59 50.6 85.0 3.26 7.25 0.90 3.64 0.95 0.23 0.84 0.14 0.77 0.14 0.39 0.05 0.32 0.05 18.9 1.02 0.77 7.55

47.9 22.4 0.75 0.64 0.09 16.2 135 0.35 35.2 12.5 2.95 18.7 2.15 16.2 63.8 3.75 3.98 0.46 28.6 5.00 20.0 1.18 8.07 0.10 3.19 221 5.13 15.5 79.7 95.9 24.3 44.9 4.92 17.6 3.17 0.25 2.80 0.42 2.29 0.46 1.42 0.19 1.27 0.19 104 0.95 0.25 13.7

17.5 1026 1.86 1.02 0.23 11.6 144 3.06 137 23.9 2.07 13.7 1.82 27.4 72.9 30.6 90.1 1.96 30.1 5.99 96.9 1.48 13.8 0.54 1.89 206 2.28 23.6 115 82.3 38.8 80.1 9.49 35.9 7.76 1.60 7.03 0.93 4.85 0.88 2.45 0.32 1.89 0.28 192 0.99 0.65 15.1

2.06 115 1.00 0.26 0.31 37.8 179 0.09 213 29.1 1.69 58.6 9.74 16.9 164 29.9 0.69 0.56 26.8 8.11 10.5 1.08 7.74 0.05 5.05 228 8.12 23.9 288 57.2 35.1 67.5 8.30 31.5 6.48 0.99 5.75 0.77 3.90 0.72 2.03 0.26 1.60 0.23 165 0.94 0.48 16.2

117 553 1.62 1.54 0.11 24.8 115 3.23 23.9 25.6 2.65 18.4 2.51 18.1 76.7 30.6 62.1 2.77 25.6 9.98 54.7 1.10 9.64 0.43 2.58 185 1.97 27.5 110 98.4 31.8 65.1 7.67 29.2 5.93 1.35 5.62 0.76 4.11 0.79 2.31 0.32 2.01 0.30 157 0.99 0.71 11.4

103 93.3 0.73 8.40 0.25 9.59 166 0.40 178 14.1 3.38 14.6 0.99 18.0 102 20.0 7.78 2.29 35.0 9.22 91.9 1.29 9.25 0.09 3.03 262 2.98 30.8 58.2 114 25.7 52.8 6.21 24.1 5.19 1.43 5.00 0.79 4.54 0.92 2.81 0.39 2.49 0.37 133 0.99 0.84 7.39

4352 3453 1.12 0.40 3.58 9.34 125 1.31 243 12.8 1.35 36.8 5.44 24.4 52.2 1439 18.5 223 22.7 59.0 69.8 1.22 10.5 0.83 7.82 161 1.82 16.7 1427 53.3 37.6 74.0 8.23 29.5 5.49 1.47 4.49 0.54 2.52 0.45 1.23 0.15 0.95 0.14 167 0.99 0.88 29.2

a b c

C-S = chlorite–smectite, K = kerolite, S = smectite, Serp = serpentine, T = talc, T-S = talc-smectite. Other abbreviations, see Table 2 and Fig. 5. Ce/Ce⁎ = 2CeN/(LaN + PrN). Eu/Eu⁎ = 2Eu /(Sm + Gd ). N

N

N

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1183-9

186

Sample #:

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

However, we note that precipitation of mixed-layer kerolite–smectite, requires not only high silica activity but also participation of reactive Al. In fact, the Al content is the main chemical difference among the members of the talc — kerolite–smectite — tri-octahedral smectite sequence (Table 3). The Al content increases with the increasing percent of smectitic layers (Fig. 8). The good correlation between Al and smectite content supports the suggestion that the tri-octahedral smectite is an endmember of this sequence. Chlorite–smectite, which is in close association with talc in the Escanaba Trough, also exhibits a positive correlation between Al content and interstratified smectitic layers (Fig. 8). However, the regression line of chlorite–smectite is much steeper than that of kerolite–smectite indicating the smectite in chlorite–smectite is more Al-rich (di-octahedral) than that in kerolite–smectite, as supported by XRD studies (Table 2; Fig. 5h, i). Submarine hydrothermal fluids contain low concentrations of dissolved Al: 4.0–5.2 μmol/l at 21°N EPR (Von Damm et al., 1985a), 17–56 μmol/l at Escanaba Trough (Von Damm et al., 2005), 2–3 μmol/l at Middle Valley (D. Butterfield, personal communication), and 0.9–7.9 μmol/l at Guaymas Basin (Von Damm et al., 1985b). The temporal and spatial variations of dissolved Al content of hydrothermal fluids could account for the different proportion of the Al-containing phase (smectitic layers) in kerolite–smectite. However, the mode of occurrence of minerals from the talc — tri-octahedral smectite sequence suggests that the Al concentration alone is not the main control on precipitation of these minerals. Our results do not support any dependence of the precipitated minerals on the type/subtype of hydrothermal system (Tables 1, 2). Our observations (Fig. 2a) and those of previous investigators (Lonsdale et al., 1980) indicate that talc occurs as colloform crusts on top of sediment, suggesting that it precipitated on the sediment surface or in open cracks near the sediment surface. The dense fracture network beneath and perpendicular to the talc-sediment contact (Fig. 2a) appears to represent a path for fluids penetrating upward through the sediment pile. Kerolite–smectite lines the inner walls of the chimney conduits, fills the interstices of cavernous massive sulfides and occurs along with chlorite, chlorite–smectite and smectite in the sediment (beneath the seafloor). Our only tri-octahedral smectite specimen (L1-86-NC-27G-77-80) occurs deep (∼ 0.8 m) in the sediment pile. We suggest the change from the more Mg-rich, Al-poor talc to the more Mgpoor, Al-rich tri-octahedral smectite we observe is related to a gradual change in the availability of Mgrich seawater. Mg is readily available on the sediment

187

surface and open cracks, is more restricted in internal parts of the sulfide/sulfate mounds and chimneys where it is supplied by advection through porous walls, and is most limited in deeper sediment horizons where it is carried downwards by penetrating seawater. Examination of the sulfate/phyllosilicate contact in chimneys shows that the minerals (Table 2, sample 2251-1-1d, serpentine + kerolite–smectite) overgrowing the sulfate are richer in Mg and less aluminous than those (sample 2251-1-1a, kerolite–smectite) overgrowing them (Table 2). Closer proximity to the chimney wall through which seawater penetrates therefore results in a phyllosilicate association richer in Mg. These observations suggest that the other main control on the precipitation of talc — kerolite–smectite — trioctahedral smectite is the Mg/Al ratio, which varies with the degree of mixing of seawater and hydrothermal fluid. The sulfide/sulfate chimney/mound shell and, to a greater extent the sediment blanket, hinder complete mixing of hydrothermal fluid and ambient seawater, resulting in a decrease in the Mg/Al ratio and precipitation of phyllosilicates poorer in Mg in the interior parts of the chimney/mound and deeper in the sediment. Sealing or clogging of chimney conduits by the serpentine + kerolite–smectite assemblage (2251-11d) limits seawater access to the hydrothermal fluid and leads to an increase in dissolved silica in the interior part of the chimney. Higher silica activity favors the formation of talc relative to serpentine (e.g., Koski et al., 1985) and results in deposition of kerolite– smectite assemblage (2251-1-1a) over the serpentine + kerolite–smectite one (2251-1-1d). The calculated formation temperatures for talc from the Escanaba Trough (287–323 °C) are in very good agreement with those calculated from previously reported δ18O data (263–313 °C) confirming that talc is a high-temperature phase (Table 5). The calculated temperatures of talc formation in Table 5 correspond to the SESCA area. They are higher than the measured vent temperature of 217 °C at the NESCA site, but are similar to the temperatures inferred from the mineral assemblages in the Escanaba massive sulfides (∼ 300 °C; Koski et al., 1994). There is a discrepancy between our δ18O value for the Fe-rich talc from the Red Seamount and that previously reported for the same sample (Table 5). However, published chemical data for this sample (Alt et al., 1987) suggest that there could be FeOOH contamination. The FeOOH impurities may have produced a net δ18O that is lower (2.8‰) than that of the pure talc separate we have analysed (4.3‰) (e.g., Poage and Chamberlain, 2001). We therefore believe

188

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Fig. 7. C1 chondrite-normalized (Sun and McDonough, 1989) REE distribution patterns for: (a) kerolite–smectite (light-grey = Grimsey samples, dark-grey = Middle Valley samples, solid triangles = S222715, solid diamonds = 7D-8b1, solid squares = 7D-27A1 GCS); (b) talc (grey = range of the Escanaba samples, solid triangles = 1183–9, solid diamonds = L1-86-NC-15D-2-3B, solid circles = L1-86-NC-15D-2-4); (c) chlorite, chlorite–smectite and smectite (grey = range of the chlorite and chlorite–smectite samples, solid squares = L1-86-NC-27G-77-80, open diamonds = L1-86-NC-15D-2-3A, open circles = L1-86-NC-15D2-4grey) samples.

that our data represent talc isotope composition better than previously published results (Alt et al., 1987) and we note that the calculated formation temperature

of this specimen falls within the range for other talc samples. We have not found pure talc in the Guaymas Basin deposits, but the formation temperature of kerolite– smectite from this site (302–306 °C) is very close to that of “talc” (272–287 °C) reported previously and within the temperature range of the venting fluids (≤315 °C, Von Damm et al., 1985b). Kerolite–smectite also appears to be a high-temperature phase (Table 5). Kerolite–smectite from the peridotite-hosted hydrothermal system at the St. Paul FZ exhibits the lowest formation temperature (226 °C) among the kerolite– smectites we have studied. This temperature substantially exceeds the maximum temperatures measured at a typical peridotite-hosted hydrothermal system like the Lost City (40–75 °C; Kelley et al., 2001), but is consistent with subsurface temperatures of 200 ± 50 °C estimated for Lost City fluids by Allen and Seyfried (2004). We speculate that the St. Paul FZ hydrothermal system is related to higher temperature (N250 °C) circulating fluids, like that of the peridotite-hosted Rainbow hydrothermal field (T ≤ 365 °C; Douville et al., 2002). We used the δ18O value of seawater (0‰) for our calculation since we had no values of the O isotope composition of the venting fluids. The only data on the O isotope composition of fluids from serpentinized peridotites are from the Conical and S. Chamorro seamounts, Mariana forearc (Mottl et al., 2003). They are 4.0 and 2.5‰, respectively. This implies that the δ18O of the St. Paul FZ fluids might have been N 0‰ which implies that the formation temperature of the precipitated kerolite–smectite could have been N 226 °C. Calculated formation temperatures for both Middle Valley chimney kerolite–smectites are consistent, giving temperatures of 278–300 °C (Table 5). For one of them (2251-1-1a, from Dead Dog Mound), the calculated temperature is very close to the measured vent fluid temperature (261–268 °C), whereas, for the other (2251–2–3, from Heineken Hollow) the temperature differs substantially from the measured vent fluid temperature of 184 °C (Percival and Ames, 1993). We speculate that, at the time when the kerolite–smectite precipitated in the Heineken Hollow chimneys, the temperature of the venting fluids was higher than that of the recent fluids. The formation temperatures of kerolite–smectite from the Grimsey Graben (319–382 °C) are the highest calculated in our sample set and similar to those previously reported by Lackschewitz et al. (2006) for kerolite–smectite from the same field (Table 5). However, these temperatures are unreasonably high considering that: (i) the measured in situ maximum vent

Table 5 Oxygen isotope composition and calculated formation temperatures for the studied samples δ18O (‰) Formation temperature (°C) Fractionation factor source

Locality

Mineral

1183-9 L1-86-NC-15D-2A L1-86-NC-15D-2-2 L1-86-NC-15D-2-1 L1-86-NC-15D-2-3B L1-86-NC-15D-2-4 2251-2-3 2251-1-1a 7D-8b1 SL 347 GC 60-80 SL 347 GC 140-160 SL 347 GC 160-180 SL 347 GC 200-220 SL 347 GC CC S2227-15 L1-86-NC-15D-3A L1-86-NC-27G-77-80 L1-86-NC-24D-7-1

Red Seamount, 21°N EPR Escanaba Trough -"-"-"-"Middle Valley, Juan de Fuca Ridge Guaymas Basin, S Trough Grimsey Graben -"-"-"-"St. Paul Fracture Zone Escanaba Trough -"-"-

Talc 4.3 Talc 4.0 Talc 3.5 Talc 3.0 Talc 3.9 Talc 4.7 Kerolite–smectite 4.2 Kerolite–smectite 4.2 Kerolite–smectite 4.4 Kerolite–smectite 1.8 Kerolite–smectite 1.5 Kerolite–smectite 2.2 Kerolite–smectite 2.5 Kerolite–smectite − 0.4 Kerolite–smectite 7.1 Talc-smectite, kerolite–smectite 3.1 Smectite 5.4 Chlorite 3.7

287 302 313 323 304 287 278-300 278-300 302-306 334 341 326 319 382 226 321 257 169-207

L1-86-NC-15D-2-4grey -"-

Chlorite, chlorite–smectite

4.3

153-192

δ18O values from the literature 1183-9 Red Seamount, 21°N EPR 7D-27 Guaymas Basin, S Trough Talc-A Guaymas Basin, N Trough GHF1 Grimsey Graben L1-86-NC-15D-2 Escanaba Trough L1-86-NC-24D-21 -"L1-86-NC-27G-77-80 -"L2-86-NC-9G-115-122 -"-

Talc Talc Talc Kerolite–smectite Talc Talc Smectite Chlorite

2.8 5.8 5.3 1.9 5.8 3.5 6.0 2.7

319 272-276 282-287 332 263 313 240 199-234

Reference

Saccocia et al. (in preparation) -"-"-"-"-"-"-"-"-"-"-"-"-"-"-"Savin and Lee (1988) Wenner and Taylor (1973); Cole and Ripley (1999) Wenner and Taylor (1973); Cole and Ripley (1999)

This study -"-"-"-"-"-"-"-"-"-"-"-"-"-"-"-"-"-

Saccocia et al. (in preparation) -"-"-"-"-"Savin and Lee (1988) Wenner and Taylor (1973); Cole and Ripley (1999)

Alt et al. (1987) Koski et al. (1985) Lonsdale et al. (1980) Lackschewitz et al. (2006) Zierenberg and Shanks (1994) -"-"-"-

-"-

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

Sample #

189

190

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

Fig. 8. Al2O3 content versus % smectite layers in chlorite (solid squares), chlorite–smectite (solid diamonds), talc (crosses), kerolite– smectite (open circles), and tri-octahedral smectite (solid triangles). Al2O3 content increases with increasing % smectite in chlorite– smectite (straight arrow) indicating that the interstratified smectite is di-octahedral. The Al2O3 content of the mineral sequence talc – kerolite–smectite – tri-octahedral smectite also increases with smectite content (curved arrow). The two kerolite–smectite datapoints with high Al2O3 correspond to a sample (L1-86-NC-15D-2-8) where smectite is mainly di-octahedral.

temperature was 251 °C (Hannington et al., 2001); and (ii) the maximum theoretically possible temperature is 250 °C according to the pressure-boiling curve of seawater at the water depth of the Grimsey hydrothermal field (400 m) (Hannington et al., 2001). The high temperatures from O isotope data for the formation of kerolite–smectites from the Grimsey Graben (319–382 °C) are puzzling because vent temperatures could not exceed the seafloor boiling temperature of 250 °C. Two explanations are possible: (1) the talc–water fractionation curve is not applicable to the kerolite–smectite samples from this site, or (2) the hydrothermal fluids that formed these samples had δ18O values less than seawater, approximately — 3‰. Our results show that application of the talc–water curve to kerolite–smectite phases from other sites yields very reasonable temperature values compared to measured vent fluid temperatures and other temperature indicators. We have therefore no reason to believe that the equation we used is not a valid approximation for kerolite–smectite. On the other hand, seafloor hydrothermal vent fluids with δ18O values significantly less than 0‰ have never been encountered (Shanks, 2001) and present-day vent fluids of Grimsey Graben have δ18O values from − 0.3 to 0.9 (Lackschewitz et al., 2006). Fluid boiling or magmatic water contributions

would produce the opposite result: vent fluids with positive δ18O values. Three processes that could have operated in the past to produce negative δ18O fluids are: (1) diagenesis of sediments (Longstaffe, 2000) at fairly low temperatures (b 150 °C) and very low water–rock ratio (b0.1); (2) seawater reaction with peridotite during serpentinization at temperatures b200 °C and water– rock ratio b 10 (Alt et al., 2007), which can produce fluids with values of − 3‰ predicted at 150 °C and water–rock ratio 1; and (3) isotopically light meteoric waters formed during the Last Glacial Maximum may have mixed with circulating seawater to create low δ18O values. How do these suggestions relate to what is known about the geologic setting of Grimsey Graben? Riedel et al. (2001) have used seismic reflection profiling to show that there is a basin immediately east of the vent field that contains N60 m of sediments that are a possible source for thermogenic hydrocarbons in the vent fluids. Perhaps low water–rock, low-temperature diagenesis occurred in this basin and produced pore fluids with negative δ18O values prior to the onset of hydrothermal activity. Once magmatism and faulting established hydrothermal convection systems, these pore fluids might have been entrained for a period of time into the circulating seawater hydrothermal system, producing mixed fluids with a net δ18O of about − 3‰. Alternatively, there could be unknown subsurface occurrences of ultramafic rocks near Grimsey Graben and serpentinization reactions could produce similar low-δ18O fluids, but this latter hypothesis is speculative at this point in time. Finally, Andrews et al. (2000) have shown that ice sheets extended from Iceland to Grimsey Island during the Last Glacial Maximum, but it is not clear how low δ 18 O waters in the ice might be incorporated into circulating hydrothermal fluids. Socalled “thermal kame” deposits are not uncommon in Yellowstone National Park as a direct result of hydrothermal venting beneath glaciers during the last ice age. Further studies, particularly on the age of the kerolite–smectite deposits and on pore fluids in the sedimentary section, are needed to better understand this apparent enigma. A smectite sample from Escanaba Trough gives a moderate δ18O temperature of formation (257 °C), which is in very good agreement with the previously reported value for the same sample (240 °C; Table 5). The calculated temperatures of formation of chlorite and chlorite–smectite (153–207 °C) are very close to that published earlier (199–234 °C; Table 5) confirming that these minerals are the lowest temperature phyllosilicates in our sample set.

V.M. Dekov et al. / Chemical Geology 247 (2008) 171–194

The high formation temperatures (generally N 250 °C) calculated for talc and kerolite–smectite indicate that they formed where there was a focused fluid discharge. This is in agreement with our interpretation of the chemical Mg/Al control on the formations of these phases, indicating that they precipitated in chimneys, in massive sulfide mounds, at the sediments surface and in open cracks in the sediment near the seawater–seafloor interface. The other end-member of this tri-octahedral Mg-phyllosilicate sequence, smectite, forms at moderate-temperature (200–250 °C), deep in the sediment blanket (∼ 0.8 m). Chlorite and chlorite–smectite, which constitute the alteration sediment matrix around the Escanaba hydrothermal mounds, are lower-temperature (150–200 °C) products of diffuse fluid discharge through the sediment around the hydrothermal conduits. The very low trace- and REE concentrations in talc and especially in kerolite–smectite support the precipitation of these minerals from hydrothermal fluids. The higher contents of these elements in chlorite and chlorite–smectite, on the other hand, imply that they have inherited them from precursor minerals and thus supports the interpretation that chlorite and chlorite–smectite are hydrothermal alteration products of hemipelagic detrital sediment (Zierenberg and Shanks, 1994). The extremely low REE contents and strong positive Eu anomaly in kerolite–smectite (Fig. 7a) corroborate the formation temperature calculations showing that this mineral precipitated from high-temperature hydrothermal fluids. Previous studies (Hannington et al., 2001; D'Orazio et al., 2004) also showed that kerolite– smectite from the Grimsey core and St. Paul FZ formed in hydrothermal vents. Because none of the hydrothermal fluids from any of the studied fields exhibit a Ce anomaly (Klinkhammer et al., 1994), the negative Ce anomaly observed in kerolite–smectite (Fig. 7a) can be attributed to the mixing of hydrothermal fluids with seawater. Kerolite–smectite from the Middle Valley chimneys and a sediment sample from the Guaymas Basin (7D-27A1 GCS) show a weak negative Eu anomaly (Fig. 7a). It is well known that the REE patterns of seafloor high-temperature end-member hydrothermal fluids are remarkably similar with a prominent positive Eu anomaly and LREE enrichment, regardless of seafloor setting (Klinkhammer et al., 1994). However, variations in the temperature and pH, duration of fluid–rock interaction, precipitation and exchange with secondary minerals, and anion complexation affect the REE patterns of vent fluids (Klinkhammer et al., 1994). Accordingly, the REE patterns of vent fluids might be expected to vary within

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a single hydrothermal field, especially in the case of more complex fields. REE data for vent fluids are not available from the studied Middle Valley chimneys and Guaymas Basin site (Table 1) and we can only speculate about the cause of the observed weak negative Eu anomaly in the kerolite– smectite samples from there (Fig. 7a). The negative Eu anomaly might imply that the parent solution from which a mineral precipitated had temperatures b250 °C upon discharge, because high-temperature hydrothermal fluids exhibit positive Eu anomalies (Michard, 1989). However, this is not the case of these 3 samples in our study because all formed above 250 °C (Table 5). In principle, this Eu anomaly reflects Eu depletion in the original hydrothermal fluid. Such Eu depletion could be caused by precipitation of minerals-sinks for Eu2+ deeper in the system. Anhydrite and barite, which form chimneys and crusts (on or within the sediment) and also appear scattered in the sediment of these fields (Koski et al., 1985; Percival and Ames, 1993), are possible sinks for Eu2+. Their subsurface precipitation is likely to cause a Eu deficiency of the residual fluid from which the phyllosilicates precipitate. The similarity of the REE patterns of our talc samples (Fig. 7b) suggests precipitation under similar conditions. The relatively low REE abundances and LREE enrichment combined with the smooth gradual HREE depletion might indicate precipitation from hydrothermal fluids. The lack or presence of a weak negative Eu anomaly is interpreted as a sign of Eu depletion in the parent high-temperature (N250 °C) fluid. Since all the talc samples have been found on top of sediment semilithified chunks or lining fissures near the sediment surface, it is probable that the precipitating fluid has been depleted in Eu as a result of subsurface precipitation of anhydrite and barite, which are widespread at the studied sites (Koski et al., 1994). The REE abundances and distribution patterns of chlorite and chlorite–smectite samples (Fig. 7c) are similar to those of the upper continental crust and continental arc turbidites (with LREE enrichment, flat HREE distributions and negative Eu anomaly) (McLennan, 1989). According to previous studies (Zierenberg and Shanks, 1994), chlorite and chlorite–smectite are hydrothermal alteration products of the background turbiditic sediment. The observed REE distribution patterns suggest that the alteration products have retained the pattern of the primary sediment. Since these are the same fluids that produced talc in open cracks and on the sediment surface, they would be depleted in Eu and would not alter the negative Eu anomaly inherited from the parent sediment.

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The low REE abundance, LREE enrichment and weak positive Eu anomaly of the REE pattern (Fig. 7c) for the only tri-octahedral smectite we studied do not discriminate between direct precipitation from hydrothermal fluid and hydrothermal alteration of primary sediment. Since the REE patterns are intermediate between those of kerolite–smectite (hydrothermal precipitate), and chlorite and chlorite–smectite (hydrothermal alteration products), we suppose that the smectite is a hydrothermal alteration product. 6. Summary Previous studies on Mg-phyllosilicates at seafloor hydrothermal fields have described talc, stevensite and saponite. Our inspection of published data reveals some of these determinations to be inaccurate. Investigation of clay minerals from six seafloor hydrothermal areas from peridotite- and basalt-hosted (sediment-free and sediment-covered) hydrothermal systems has revealed a sequence of tri-octahedral Mg-phyllosilicates ranging from pure talc and Fe-rich talc, through kerolite-rich kerolite–smectite to smectite-rich kerolite–smectite and tri-octahedral smectite. The most common mineral is mixed-layer kerolite–smectite showing a series of interstratification with 5 to 85% smectitic layers. The smectite interstratified with kerolite is predominantly tri-octahedral. The degree of crystal perfection of the clay sequence decreases from talc to kerolite–smectite, with lower crystalline perfection as the proportion of smectite layers in kerolite–smectite increases. Our investigations did not reveal any dependence of the precipitated minerals on the type/subtype of hydrothermal system. Talc and kerolite–smectite precipitated in the chimneys, massive sulfide mounds, at the sediment surface and in open cracks in the sediment near the seawater–seafloor interface from focused discharge of high-temperature fluids (N 250 °C). The other endmember of this tri-octahedral Mg-phyllosilicate sequence, smectite, is a moderate-temperature (200–250 °C) phase that formed deep in the sediment. Chlorite and chlorite– smectite make up the altered sediment matrix around the hydrothermal mounds. These minerals formed at low temperature (150–200 °C) as products of diffuse fluid discharge through the sediment around the hydrothermal conduits. The chemical controls on mineral precipitation in the talc–smectite sequence are silica activity and Mg/Al ratio which, in turn, depend on the degree of mixing of seawater and hydrothermal fluid. Higher silica activity and Mg/Al ratio favor the precipitation of talc relative to tri-octahedral smectite. The vent structures and sediment blanket hamper complete mixing of hydrothermal fluid

and ambient seawater, producing a decrease of the Mg/Al ratio and precipitation of Mg phyllosilicates with lower Mg content in the interior parts of the chimneys and deeper in the sediment (i.e., kerolite–smectite, smectite). Talc and kerolite–smectite have very low concentrations of trace elements and REE. Although they are high-temperature phases, some of them exhibit negative or no Eu anomaly, which suggests Eu depletion in the parent hydrothermal fluid, probably caused by precipitation of minerals sinks for Eu2+ deeper in the system (anhydrite, barite). The REE abundances and distribution patterns of chlorite and chlorite–smectite imply that they are hydrothermal alteration products of the background turbiditic sediment. Acknowledgements We are very grateful to J.C. Alt (University of Michigan, USA), D.E. Ames (Geological Survey of Canada), M. D'Orazio (University of Pisa, Italy) and T. Kuhn (IFM-GEOMAR, Germany) who provided the samples from the Red Seamount, Middle Valley, St. Paul FZ and Grimsey Graben, respectively. We thank also D. Butterfield (University of Washington, USA) for the dissolved Al data from the Middle Valley fluids. This study was supported by a Royal Society Grant (2005/ R1-JP). V. Dekov was partly funded by the Alexander von Humboldt Foundation (Rueckehr Stipendium). Thanks go to D. Rickard, G.P. Glasby and the anonymous reviewer for their valuable suggestions, which substantially improved this paper. References Allen, D.E., Seyfried Jr., W.E., 2004. Serpentinization and heat generation; constraints from Lost City and Rainbow hydrothermal systems. Geochimica et Cosmochimica Acta 68 (6), 1347–1354. Alt, J.C., Lonsdale, P., Haymon, R., Muehlenbachs, K., 1987. Hydrothermal sulphide and oxide deposits on seamounts near 21°N, East Pacific Rise. Geological Society of America Bulletin 98 (2), 157–168. Alt, J.C., Shanks III, W.C., Bach, W., Paulick, H., Garrido, C.J., Beaudoin, G., 2007. Hydrothermal alteration and microbial sulfate reduction in peridotite and gabbro exposed by detachment faulting at the Mid-Atlantic Ridge, 15°20′N (ODP Leg 209): A sulfur and oxygen isotope study. Geochemistry, Geophysics, Geosystems 8 (8), Q08002. doi:10.1029/2007GC001617. Ames, D.E., Franklin, J.M., Hannington, M.D., 1993. Mineralogy and geochemistry of active and inactive chimneys and massive sulphide, Middle Valley, northern Juan de Fuca Ridge: an evolving hydrothermal system. Canadian Mineralogist 31 (4), 997–1024. Andrews, J.T., Hardardóttir, J., Helgadóttir, G., Jennings, A.E., Geirsdóttir, Á., Sveinbjörnsdóttir, Á.E., Schoolfield, S., Kristjánsdóttir, G.B., Smith, L.M., Thors, K., Syvitski, J.P.M., 2000. The N and W Iceland Shelf: insights into Last Glacial Maximum ice

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