Modern Ocean Floor Processes and the Geological Record (Geological Society Special Publication)

Modern Ocean Floor Processes and the Geological Record

Geological Society Special Publications Series Editors

A. J. FLEET R. E. HOLDSWORTH A. C. MORTON M. S. STOKER

References to this volume It is recommended that reference to all or part of this book should be made in one of the following ways:

MILLS,R. A. & HARRISON,K. (eds) 1998. Modern Ocean Floor Processes and the Geological Record. Geological Society, London, Special Publications, 148.

COLLIER,J. S. • SINGH, S. C. 1998. A seismic inversion study of the axial magma chamber reflector beneath the East Pacific Rise near 10°N. In." MILLS, R. A. 8,: HARRISON, K. (eds) 1998. Modern Ocean Floor Processes and the Geological Record. Geological Society, London, Special Publications, 148, 17-28.

G E O L O G I C A L SOCIETY SPECIAL P U B L I C A T I O N NO. 148

Modern Ocean Floor Processes and the Geological Record EDITED BY

R. A. MILLS School of Ocean and Earth Science, University of Southampton, UK and

K. H A R R I S O N BRIDGE Programme Manager, University of Leeds, UK

1998 Published by The Geological Society London

THE GEOLOGICAL

SOCIETY

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Contents Preface

vii

GR~kCIA,E., PARSON,L. M. & BIDEAU,D. Volcano-tectonic variability along segments of the Mid-Atlantic Ridge between Azores platform and Hayes fracture zone: evidence from submersible and high-resolution sidescan sonar data

COLLIER,J. S. & SINGH, S. C. A seismic inversion study of the axial magma chamber reflector beneath the East Pacific Rise near 10°N

17

ALLERTON, S. • MACLEOD, C. J. Fault-controlled magma transport through the mantle lithosphere at slow-spreading ridges

29

DILEK. Y. & TnY, P. Structure, petrology and seafloor spreading tectonics of the Kizildag Ophiolite, Turkey

43

MINSHULL,T. A., MULLER,M. R., ROBINSON,C. J., WHITE, R. S. & BICKLE,M. J. Is the oceanic Moho a serpentinization front?

71

TEAGLE,D. A. H., AET, J. C. & HALLIDAY,A. N. Tracing the evolution of hydrothermal fluids in the upper oceanic crust: Sr-isotopic constraints from DSDP/ODP Holes 504B and 896A

81

HUNTER, A. G. & ODP LEG 168 SCIENTIFICPARTY. Petrological investigations of low temperature hydrothermal alteration of the upper crust, Juan de Fuca Ridge, ODP Leg 168

99

BICKLE, M. J., TEAGLE,D. A. H., BEYNON,J. & CHAPMAN,H. J. The structure and controls on fluid-rock interactions in ocean ridge hydrothermal systems: constraints from the Troodos ophiolite

127

WELLS, D. M., MILLS, R. A. & ROBERTS, S. Rare earth element mobility in a mineralized alteration pipe within the Troodos ophiolite, Cyprus

153

JAMES, R. H., DUCKWORTH,R. C., PALMER,M. R. & THE ODP LEG 169 SHIPBOARD SCIENTIFICPARTY. Drilling of sediment-hosted massive sulphide deposits at the Middle Valley and Escanaba Trough spreading centres: ODP Leg 169

177

GOULDING, H. C., MILLS, R. A. & NESBITT, R. W. Precipitation of hydrothermal sediments on the active TAG mound: implications for ochre formation

201

ROBERTSON,A. & DEGNAN,P. Significance of modern and ancient oceanic Mn-rich hydrothermal sediments, exemplified by Jurassic Mn-cherts from Southern Greece

217

HERRINGTON, R. J., MASLENNIKOV,V. V., SPIRO, B., ZAYKOV,V. V. & LITTLE, C. T. S. Ancient vent chimney structures in the Silurian massive sulphides of the Urals

241

LITTLE, C. T. S., HERRINGTON, R. J., MASLENNIKOV,V. V. & ZAYKOV,V. V. The fossil record of hydrothermal vent communities

259

MCARTHUR, A. G. & TUNNICLIFFE,V. Relics and antiquity revisited in the modern vent fauna

271

Index

293

The meeting that led to this volume was sponsored by the following three bodies whose support is gratefully acknowledged.

Marine Studies Group Geological Society of London The Marine Studies Group is a specialist group of the Geological Society. It provides a focus for marine geoscientists and promotes the exchange of ideas within the UK community. It exists to develop links between researchers of different disciplines within the marine geoscience field, to strengthen contact with land-based specialist groups and to maintain a high profile for geological progress in the field amongst technologists and industrialists. Conferences sponsored each year include those aimed specifically at undergraduate and postgraduate research topics. The group has strong links with non-Society organizations such as The Challenger Society and BRIDGE, and welcomes contact with other marine-minded bodies. The Marine Studies Group can be contacted via the Geological Society.

British Mid-Ocean Ridge Initiative (BRIDGE) Natural Environment Research Council BRIDGE is a thematic research programme of the UK's Natural Environment Research Council (NERC). Running from 1993 to 1999, BRIDGE investigates the creation of new ocean crust at midocean ridges. The volcanic mid-ocean ridge system is unusually compact and provides an opportunity to understand the oceanographic, geological, chemical and biological aspects of a complete environment. BRIDGE investigates the geological setting of the ridge: the geochemistry of vent fluids and hydrothermal mounds; and ways in which biological communities survive in this apparently hostile environment. Research ranges from regional scale studies mapping unexplored seafloor to microscopic and chemical analyses at individual vent sites. BRIDGE develops novel deep ocean technologies for deployment from surface vessels and manned submersibles, and funds experimental research into the mechanical and chemical nature of the rocks and underlying crust of these active volcanic zones. Contact: BRIDGE Programme Manager, School of Earth Sciences, University of Leeds, Leeds LS2 9JT, UK or BRIDGE Programme, Natural Environment Research Council, Swindon SN2 1EU, UK.

The Challenger Society for Marine Science Founded in 1903 the Challenger Society's objectives are: to advance the study of Marine Science through research and education; to disseminate knowledge of Marine Science with a view to encouraging a wider interest in the study of the seas and an awareness of the need for their proper management; and to contribute to public debate on the development of Marine Science. The Society aims to achieve these objectives by holding regular scientific meetings covering all aspects of Marine Science, by supporting specialist groups to provide a forum for discussion, and by publishing a range of documents dealing with aspects of Marine Science and the Society's programme of meetings. The Society's journal, Ocean Challenge, is published three times a year. For details of membership or other Society matters contact: The Challenger Society for Marine Science, Southampton Oceanography Centre (251/20), Southampton SO14 3ZH, UK. Other Geological Society Special Publications sponsored by these three organizations are: No. 87 Parson, L. M., Walker, C. L. & Dixon, D. R. (eds) 1995. Hydrothermal Vents and Processes. No. 118 MacLeod, C. J., Tyler, P. A. & Walker, C. L. (eds) 1996. Tectonic, Magmatic, Hydrothermal

and Biological Segmentation of Mid-Ocean Ridges.

Preface

A multidisciplinary volume bringing together studies of the geologically active sea floor and the ancient record of such activity is long overdue. With the recent achievements of the BRIDGE and RIDGE programmes, and the successful drilling of oceanic crust and areas of mineralisation, it is timely to test the new models against the wealth of data for similar systems found throughout the geological record. As our aim is to synthesise knowledge of active and preserved sea floor, we have juxtaposed papers based on the ideas and hypotheses considered, rather than geographical or geological order. Thus the reader is taken from the structure of mid-ocean ridges, through crustal alteration and mineralisation processes to a consideration of faunal communities. The first five papers all concern crustal configuration. Gr~eia et al. combine evidence from submersible and high resolution side scan sonar to study the relationship between volcanism and tectonism south of the Azores. While the scales of these modes of observation are not wholly compatible, the records do allow segment-scale inferences to be made on the crustal magma system. Collier & Singh have carried out a novel seismic inversion technique to allow imaging of a melt lens on the East Pacific Rise and limits to be put on the amount of melt present. Their results also have implications for the segmentation of the crustal magma system, though here at a fast spreading ridge. Allerton & MaeLeod make inferences of magma transport from Mid-Atlantic Ridge topography and conclude that melt may be channelled through the lithospheric mantle by segment-end faults. They test this hypothesis in the Carrick Luz shear zone of the Lizard Ophiolite, Cornwall, UK where gabbro mylonitic shear zones are observed. Dilek & Thy describe detailed observations of the structure and petrology of the Kizildag Ophiolite, Turkey in the light of observations from the Troodos Ophiolite, Cyprus. This section ends with Minshuli et al. revisiting a question posed by Hess in the 1960s - ' i s the oceanic Moho a serpentinisation front?' They argue succinctly and convincingly that this hypothesis holds in areas of thin crust at fracture zones. Crustal alteration is addressed by the next four papers. All the convincing evidence from modern seafloor settings comes from Ocean Drilling Program results. Teagle et al. record fluid evolution through the crust from secondary mineral strontium isotope composition during recharge and discharge at the site of ODP hole 504B in 6 Ma old crust on the East Pacific Rise. Such records are vital to our understanding of crustal evolution and fluid fluxes. Hunter et al. record the petrology of low-temperature alteration of the upper crust on the Juan de Fuca Ridge, again providing new insights into fluid flow through crust of medium spreading rate. BieMe et al. and Wells et al. describe studies from the Troodos Ophiolite. Bickle et al. use strontium isotope systematics to study the extent of crustal alteration and highlight the difficulty in reconciling the degree of alteration of Troodos sheeted dykes with the inferred fluid fluxes over time. Wells et al. study alteration on a more localised scale and use novel laser 1CPMS measurements to identify the timing of reaction in an alteration pipe beneath a mineralisation zone, again highlighting the importance of the low temperature overprint on the geochemical budgets for Troodos. Both high and low temperature mineralisation at mid-ocean ridges are addressed in the next four papers. Ocean Drilling Program sampling of sediment-hosted massive sulphides on the Juan de Fuca and Gorda Ridges is discussed by James et al. The authors couple sulphide petrological studies with pore fluid analyses to describe the mode of formation for the two deposits drilled. The interaction of fluids with mineralised material is addressed by Goulding et al. in the study of ochreous sediments from the active TAG hydrothermal mound. Basalt replacement and redox zonation of metals are the important controls on the geochemistry of such sediments. Robertson & Degnan review manganeserich hydrothermal sediment occurrences from the modern oceans and use models of formation to infer the origin of Jurassic Mn-cherts from Southern Greece. Moving back through the geological record, Herrington et al. describe superbly preserved vent chimney structures from Silurian deposits in the Urals, Russia. Again, models of modern chimney formation explain the features observed in these ancient deposits.

viii

PREFACE

The final two papers in the volume bring together biological studies from modern sea floor systems and the ancient palaeontological record to focus on community structure. Little et al. describe fossiliferous deposits from the Urals and test the refuge hypothesis posed from studies of modern systems. MeArthur & Tunnieliffe also address this hypothesis and surmise that hydrothermal vent environments may provide protection from certain causes of extinction through geological time. The challenges to scientists now are to take the ideas presented forward, to test the hypotheses constructed and to integrate more thoroughly studies of the deep ocean with studies of the deep past. The focus of such studies will necessarily move away from mid-ocean ridges, not the best analogues of ophiolite settings, and move to back-arc basins and areas of crustal convergence. Recent discoveries of high-grade, base metal-rich deposits in the Manus Basin demonstrate that our understanding of processes from mid-ocean ridges is applicable to other oceanic environments. Discovery of mineral deposits within Exclusive Economic Zone limits and at relatively shallow depths brings the potential for exploitation much closer. Before commercial factors overwhelm scientific objectives, the community should apply its new found expertise and common interests to these areas. This volume arises from a meeting held in May 1997 under the sponsorship of the Geological Society's Marine Studies Group, BRIDGE and the Challenger Society for Marine Science. The enthusiastic response from the community, both in the presentation of new ideas and the interest of the audience, demonstrated the need for a volume recording the output of the meeting. Over the following months, authors have fleshed out the ideas presented in London, argued the detail during the review process and here present papers that focus on the interface between modern and ancient sea floor processes. We are indebted to the referees who gave their time to review the manuscripts and who made so many valuable observations. Thanks go to Jeff Alt (Michigan), Tim Barrett (British Columbia), Derek Briggs (Bristol), Joe Cann (Leeds), Gail Christeson (Texas), Bob Derrick (Woods Hole), Rowena Duckworth (James Cook), Aline Fiala-Medioni (Banyuls-sur-Mer), Andy Fleet (Natural History Museum, London), Peter Floyd (Keele), Dan Fornari (Woods Hole), Chris German (SOC), Kathy Gillis (Victoria), Geoff Glasby (cosmopolitan), Mark Hannington (Geological Survey of Canada), Greg Harper (Albany), Bob Hessler (Scripps), Jose Honnorez (Strasbourg), John Hudson (Leicester), Susan Humphris (Woods Hole), Pamela Kempton (NERC Isotope Geosciences Laboratory), Mike Kendall (Leeds), David Needham (IFREMER), Bob Nesbitt (Southampton), William Newman (Scripps), Adolph Nicolas (Montpellier), Hazel Prichard (Wales), Steve Roberts (Southampton), Alastair Robertson (Edinburgh), Peter Rona (Rutgers), Martin Sinha (Cambridge), Norm Sleep (Stanford), Damon Teagle (Michigan), Meg Tivey (Woods Hole), Doug Toomey (Oregon), Paul Tyler (Southampton), Eva Valsami-Jones (Natural History Museum, London), Cindy Van Dover (Alaska), Soterios Varnavas (Patras), Bob Whitmarsh (Southampton) and Bob Zierenberg (California), with apologies to anyone we have missed. Rachel Mills Keith Harrison

Contents Preface

vii

GR~kCIA,E., PARSON,L. M. & BIDEAU,D. Volcano-tectonic variability along segments of the Mid-Atlantic Ridge between Azores platform and Hayes fracture zone: evidence from submersible and high-resolution sidescan sonar data

COLLIER,J. S. & SINGH, S. C. A seismic inversion study of the axial magma chamber reflector beneath the East Pacific Rise near 10°N

17

ALLERTON, S. • MACLEOD, C. J. Fault-controlled magma transport through the mantle lithosphere at slow-spreading ridges

29

DILEK. Y. & TnY, P. Structure, petrology and seafloor spreading tectonics of the Kizildag Ophiolite, Turkey

43

MINSHULL,T. A., MULLER,M. R., ROBINSON,C. J., WHITE, R. S. & BICKLE,M. J. Is the oceanic Moho a serpentinization front?

71

TEAGLE,D. A. H., AET, J. C. & HALLIDAY,A. N. Tracing the evolution of hydrothermal fluids in the upper oceanic crust: Sr-isotopic constraints from DSDP/ODP Holes 504B and 896A

81

HUNTER, A. G. & ODP LEG 168 SCIENTIFICPARTY. Petrological investigations of low temperature hydrothermal alteration of the upper crust, Juan de Fuca Ridge, ODP Leg 168

99

BICKLE, M. J., TEAGLE,D. A. H., BEYNON,J. & CHAPMAN,H. J. The structure and controls on fluid-rock interactions in ocean ridge hydrothermal systems: constraints from the Troodos ophiolite

127

WELLS, D. M., MILLS, R. A. & ROBERTS, S. Rare earth element mobility in a mineralized alteration pipe within the Troodos ophiolite, Cyprus

153

JAMES, R. H., DUCKWORTH,R. C., PALMER,M. R. & THE ODP LEG 169 SHIPBOARD SCIENTIFICPARTY. Drilling of sediment-hosted massive sulphide deposits at the Middle Valley and Escanaba Trough spreading centres: ODP Leg 169

177

GOULDING, H. C., MILLS, R. A. & NESBITT, R. W. Precipitation of hydrothermal sediments on the active TAG mound: implications for ochre formation

201

ROBERTSON,A. & DEGNAN,P. Significance of modern and ancient oceanic Mn-rich hydrothermal sediments, exemplified by Jurassic Mn-cherts from Southern Greece

217

HERRINGTON, R. J., MASLENNIKOV,V. V., SPIRO, B., ZAYKOV,V. V. & LITTLE, C. T. S. Ancient vent chimney structures in the Silurian massive sulphides of the Urals

241

LITTLE, C. T. S., HERRINGTON, R. J., MASLENNIKOV,V. V. & ZAYKOV,V. V. The fossil record of hydrothermal vent communities

259

MCARTHUR, A. G. & TUNNICLIFFE,V. Relics and antiquity revisited in the modern vent fauna

271

Index

293

Volcano-tectonic variability along segments of the Mid-Atlantic Ridge between Azores platform and Hayes fracture zone: evidence from submersible and high-resolution sidescan sonar data EULALIA

G R , ~ C I A 1, L I N D S A Y

M . P A R S O N 1, D A N I E L

BIDEAU 2 & ROGER

HEKINIAN 2

1Challenger Division jor Seafloor Processes, Southampton Oceanography Centre, Empress Dock, Southampton S 0 1 4 3 Z H , UK 2 D R O / G M , IFRE~I4ER-Centre de Brest, B P 70, 29280 Plouzand, France Abstract: Three contrasting segments of the Mid-Atlantic Ridge (MAR) between the Azores platform and Hayes fractme zones have been studied to determine variations in the magmato-tectonic processes between segment centre and segment ends. Two of the segments (OH1 and OH3) were surveyed with the submersible Nautile and the third (Lucky Strike or PO1), with high resolution deep-tow side-scan sonar (TOBI). Segment OHI is a long and robust segment with a wide and shallow axial valley floor. Submersible observations support the inference that this segment is characterized by intense volcanic activity concentrated principally at the segment centre. Segment OH3, with a narrow and deep rift valley, is dominated by tectonic features at the segment centre, whereas the most recent volcanic constructions are found to outcrop in its southern part. Lucky Strike contains a shallow centrally located volcanic platform but tectonic features predominate all along the segment. The only evidence of fresh volcanism are a lava lake and fresh sheetflows located at the middle and northern part of the platform, respectively. Each of the three segments show 'bull's eye' gravity lows suggesting a focused mantle diapiric source.

Along-axis volcano-tectonic variability has been commonly observed on the Mid-Atlantic Ridge (MAR) (e.g. Karson et al. 1987; Semp6r6 et al. 1993), and seems to correspond to the surface expression of the segmentation processes occurring at depth (e.g. Lin et al. 1990). However, the patterns of along-axis distribution of volcanic and tectonic features seem to be unique to each individual segment, and can be tentatively related to different stages of segment evolution. In this paper we present the main results of a detailed study of three second-order segments of the M A R located between the Azores Triple Junction (38°30"N, 30°00'W) and Hayes fracture zone (33°30'N, 37°50'W) (Fig. 1). Two of the segments, OH1 and OH3, are located between the Oceanographer and Hayes fracture zones and were surveyed using the manned deep submersible Nautile (Bideau et al. 1996a). The third segment is Lucky Strike (POD, located between the Pico offset and Oceanographer fracture zone, which was explored using the deep-towed high resolution side-scan sonar TOBI (Towed Ocean Bottom Instrument; Flewellen et al. 1993) (German et al. 1996). The segments OH1 and OH3 are only 65 km apart but display contrasting morphologies and gravity signatures (Detrick et al. 1995). The

northern segment, OH1, is strongly magmatic with a narrow and shallow active axial floor to the east and robust off-axis seamount chains. The segment is associated with one of the largest mantle Bouguer anomalies (MBA) of the northern M A R (Detrick et al. 1997). The southern segment (OH3) is shorter, has a deep, U-shaped axial valley and the gravity low is smaller and non-central. In order to assess along- and across-strike variations in axial processes, a series of in-situ geological sections from segment centre to segment ends have been completed during the O C E A N A U T cruise (RV Nadir and submersible Nautile; Bideau et al. 1996a, b) (Fig. 2a, b). Lucky Strike segment (PO1) seems to be of a form intermediate to the above examples. It is characterized by a shallow and wide central volcanic platform, although tectonic features are commonly observed throughout the segment. The segment is also associated with a large negative MBA low suggesting focused magmatism. The objective of this paper is to compare the style of crustal accretion of these 3 segments with the F A M O U S segment (PO3) also located along this part of the M A R but showing different volcano-tectonic characteristics and thus, a probable distinct style of accretion. Several regional studies integrated with inter-

GRACIA, E., PARSON,L. M. & B1DEAU,D. 1998. Volcano-tectonic variability along segments of the Mid-Atlantic Ridge between Azores platform and Hayes fracture zone: evidence from submersible and high-resolution sidescan sonar data. In." MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Recor& Geological Society, London, Special Publications, 148, 1-15

2

E. GRACIA E T AL. 60°W 50°N

50°W f

40°W I

NORTH AMERICAN PLATE

40°N

30°W }

....

20°W I

10°W I

50°N

EURASIAN PLATE 4ov F.Z.

Lucky Strike

mo~)

40°N

(PO3)~ OH1~vOcea Famous

~=JO ~

F.Z.

Hayes F.Z.

30°N

30°N

"b~"~e Kane F.Z. 20°N

, 60°W

50°W

~'~ ~

AFRICAN PLATE

,

,

40°W

30°W

/ , 20°W

r

20°N

10°W

Fig. 1. Location map of the three studied segments along the Mid-Atlantic Ridge between the Azores Triple Junction and Hayes fracture zone. From north to south: Lucky Strike segment (POI), segment OH1 and segment OH3. The FAMOUS segment (PO3), discussed at the end of the paper, is also located.

national scientific programs, such as F A M O U S , AMAR, FARA-InterRidge and M A R F L U X , have been completed along the more than 500 km long section of the M A R between Pico offset and Hayes fracture zone (Fig.l). The whole area was surveyed in detail in 1991 during the S I G M A cruise of the RV l ' A t a l a n t e (Needham et al. 1992). The data consist of full coverage of Simrad EM-12D swath bathymetry and acoustic backscattering as well as geophysical data (Detrick et al. 1995). During the subsequent F A Z A R cruise, a systematic rock and water sampling program was carried out (every 5 to 10 km) with stations along the M A R between 41°N and 32°N (Charlou et al. 1993; Klinkhammer et al. 1993; Langmuir et al. 1993a). In addition, the axis between 38°N to 36°N was systematically surveyed during the H E A T cruise in 1994 with the deep-towed sidescan sonar TOBI (German et al. 1996). TOBI sidescan data provide high resolution acoustic backscatter data at a spatial resolution of a few metres. Pixel size is approximately 6 × 2 m in near range, approximating to the field of view possible by direct observation from a manned submersible or Remotely Operated Vehicle (ROV). Detailed submersible studies along this portion of the M A R comprise in situ geological ( A R C Y A N A 1975; Ballard et al. 1975; Crane & Ballard 1981) and hydrothermal activity exploration and sampling (Langmuir et al. 1993b; Fouquet et al. 1994).

The M A R between Pico offset and Hayes fracture zone is composed of 13 second-order segments separated by small left-lateral nontransform discontinuities (NTDs). The segments are referred to as PO1 to PO8 (in the Pico to Oceanographer section) and OH1 to OH5 (in the Oceanographer to Hayes section) after the proposed, informal nomenclature of Detrick et al. (1995). The general trend of the ridge is 010 ° and the average full spreading rate is 22 mm a -1 (DeMets et al. 1990). From north to south, the studied segments are described as follows. The Lucky Strike segment (PO1) maintains a constant width of 11 12km throughout its 60 km length, from 37°35'N to 37°05'N, and is bounded to both north and south by broad NTDs (,Fig. 2a). The segment forms a parallelsided rift flanked by walls shallowing by 1000 m from an average axial depth of 2200 m. The axial depth of the segment, decreases gradually from 3200 m and 2925 m at the intersections with the N T D s to the north and south, respectively. However, the dominant bathymetric feature is the centrally located shallow platform area occupying the axial valley between 2100m and 1600 m water depth (Fig. 2a). Its gently sloping upper surface is surmounted with a number of small volcanoes, three of which are clustered centrally at 37°18'N-32°17'W, and others which coalesce to form a constructional volcanic ridge to the west. Within this central platform, the first lava-lake observed in the M A R was reported

32'~10W

a) LuckyStrike (PO1)

32~'20W

3~t0W

36°30W

36~2~W

b) Segment OH1

36°40W

36°40W

36~20W

36°20W

36 ~'t OW

36"10W

37~50W

37°50W

c) Segment OH3

38°W

38°W

)T=40W

37~40W

37°30W

37~30W

~50N

34~N

Fig. 2. Simplified bathymetric maps of the three studied segments. The coverage of TOBI mosaic and location of the Nautile dives referred to in the text are also included. 9riginal bathymetric data from which these maps have been drawn were acquired using the Simrad EM-12D swath bathymetry system during the SIGMA cruise iNeedham et al. 1992). Contour interval is 200 m for (a) and 250 m for (b) and (e). E.T.: Eastern Trough: W.T.: Western Trough.

37~t0N

37~0N

~7~30N

32~20W

4

E. GRACIA E T AL.

(Fouquet et al. 1995) around which sulphide deposits and active hydrothermal vents have been mapped (Langmuir et al. 1993b; Fouquet et al. 1994). Segment OHI is approximately 90 km long, extending from 35°15'N to 34°32'N. The segment lies immediately south of Oceanographer fracture zone and is bounded to the south by a 30kin wide NTD centred at 34°30'N which offsets the MAR left-laterally by 30km. The axial zone has an hourglass-shaped rift-valley in plan view (Fig. 2b). The segment centre (34°52'N) is both very narrow (4km) and shallow (2200m). The valley deepens (to 4100m and 3300m) and widens (to 24kin and 12.5kin) towards the northern and southern segment ends, respectively. Between 34°50'N and 35°N and central to the segment there is a median ridge (Fig. 2b). The ridge is 25 km long, up to 5 km wide, shallows to 1675 and separates the rift valley floor into two troughs: Western and Eastern (Fig. 2b). The present-day axial valley is located on the Eastern Trough (Bideau et al. 1996a). Two chains of large and shallow off-axis seamounts normal to the axis can be observed at each flank of the ridge (Fig. 2b). Segment OH3 is 46kin long~ extending from 34°07'N to 33°43'N, and its axial depth profile shoals from a 3850m deep southern nodal basin to an axial bathymetric minimum of 3000 m at 33°57'N. The segment is bounded by broad NTDs up to 25 km wide which offset the ridge axis left-laterally by 30-35km (Fig. 2c). The axial zone shows a typical rift-valley morphology in cross-section. The rift valley is defined by a linear, parallel and wide outer valley cut by an inner valley delimited by the 3000m isobath. The inner valley also shows an hour-glass shape as in segment OH1, increasing from 2 km wide at 33°58'N to 7-9km at the segment ends (Grficia et al. 1997b) (Fig. 2c). The MBA pattern obtained from gravity data in the Lucky Strike (PO1) segment shows one of the largest peak to trough anomalies ( - 1 9 m G a l ) observed in the section between Pico offset and Oceanographer fracture zone (Detrick et al. 1995). Differences in crustal thickness of at least 2km are required between the segment centre and ends in order to explain the observed MBA (Derrick et al. 1995). Segment OH I shows a large ( - 4 0 r e G a l ) characteristic 'bull's eye' axial gravity low over the segment midpoint (Detrick et al. 1995). In contrast, the MBA pattern obtained along OH3 shows a smaller ( < - 2 0 m G a l ) 'bull's eye' gravity low which is not located at the mid-point of the segment, but is displaced 10 km towards the north. However, on both OH

segments a difference in crustal thickness between the NTDs and the segment centre of more than 5 km is required to explain the magnitude of the residual MBA observed (Detrick et al. 1995).

Results Here we report results from 12 N a u t i l e dives to two segments, OH1 and OH3, and compare these with published data and TOBI data from segment PO1 (Lucky Strike). Segment

OHI:

dive results

A program of seven dives was designed to explore the fine-scale variability of new tectonism and new volcanism along and across the segment OH1 (Fig. 2b). In this paper, we focus on the tectono-volcanic variations along the rift valley floor and axial valley walls. The topography at the central axial valley is elevated, smooth and remarkably flat, corresponding to the most recent volcanic zone of the whole segment (Fig. 3a). The area, explored during dives OT1, OT2 and OT5 (Figs 2b, 3a), is dominated by fresh sheet-flow lavas with flat, ropy and wrinkled textures, as well as rugged flows composed of broken wrinkled and brecciated lavas. Sparsely distributed, lightly sedimented pillowed cones (up to 20 m high) are isolated and surrounded by the recently erupted fluid flows. Drained lobate lavas and foundered lava lakes (Fig. 3a) are also commonly observed at this part of the segment, but their structures are of smaller scale than those observed in the East Pacific Rise (e.g. Francheteau et al. 1979). Northward and southward of the segment centre (Fig. 4aI), volcanic activity decreases and tectonic features, represented by fissure swarms, open cracks and fault scarps with fresh basaltic rubble are commonly observed. The northern part of the segment (dive OT17) is characterized by a rough and fractured topography composed by a succession of horst and graben structures (Fig. 4aII). Sediment input and cover increases significantly towards the northern end of the median ridge, and is locally preserved, for example, as a 0.5m thick layer of lithified sediment interbedded between pillow flows observed along dive OT17. The eastern flank of the median ridge is irregular and steep, characterized by recent tectonic activity, fault scarps, exposed sections of pillow lava and striated breccia with fresh talus at the foot (Fig. 3a). At the mid slope (1900m depth), extinct hydrothermal chimneys and white hydrothermal deposits are observed.

VOLCANO-TECTONIC VARIABILITY ALONG THE MAR WNW

ESE

0]

MedianRldge

1000

0T15 ..

5

,

07"1 "o

..

3000

,o"

AISX,

OT5

| _

''°

.,

Western Trough

40000

5 ................ 10~ ..........

1'5

a) Segment OH1 .c: 0/ C~

~ooo]

L•

ooo

AXIS

o,,, •

/

o,,v

*

..

.....

3000

4000

0 5 b) Segment OH3

10

15

Distance in kilometres (shem.flows, lavalake~}

J

fault

V,E,:2,5

Fig. 3. Schematic geological sections across the centre of (a) segment OH1 and (b) segment OH3. Nautile dives are located. Note the morphological and lithological contrast between both rift valley floors. V.E.: Vertical Exaggeration.

At the top of the ridge, hyaloclastites and vesicular pillow lavas are commonly observed. The western flank of the ridge is smooth with a gentle slope and comprises three main steps separated by wide sedimented platforms. The Western Trough is inactive and characterized by a rough faulted topography with some sedimentary cover (Fig. 3a) Both rift walls at the segment centre (Fig. 2b) consist of a succession of small fault scarps with narrow crests (Fig. 3a). The throw of the faults increases upwall and striated breccias are exposed along their surfaces. At the eastern wall, along the flank of an off-axis seamount (Fig. 3a), slabs of compacted hyaloclastite out-

crop within highly vesicular pillow lavas.

Segment OH3. clive results A programme of five dives was designed to explore the segment OH3 (Fig. 2c), along the axial valley floor and across the rift valley walls at the segment centre. The northern and central inner valley floor (dives OT10 and OT9, respectively) is narrow and characterized by a smooth topography dipping gently towards the north (Fig. 4bI). Most of the seafloor surveyed is represented by sheet-flows, and lobate lavas surrounding collapsed lava lakes, all of them draped by an

,;

o

0

0

3

0

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4

~o

NNE

NNE

0

s

0

ssw

~

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2400

2000

0

1

1

2

2

2

3

3 N

Distance

4

4

2600

2400

2800

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~

2800

OT2

2400

1

2'0

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2400 ~

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0T19

2000 ] SSW

in

o

II)

0

3400

,

GEOLOGY

s

0

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2800~

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S~W

1

1

,

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. . . . . .

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I I

32013 ~

kilometres

b)

'

2

2

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Segment

,

3

3

I

o1"1

4

4

3400

~°

NNE280O

,",~VE :I • 3200

.......

--it

, 'o~[

2o

..............

,', .........

3o

OT10

OH3

07"9 ~

Fig. 4. Bathymetric and schematic geological sections along the axis of (a) segment OH I and (b) segment OH3. The most recent volcanism in OH1 is found at the segment centre, whereas in segment OH3 it is displaced towards the south. V.E.: Vertical Exaggeration.

~L Cb

¢::

E

a)

p.,

> m

>.

C~

VOLCANO-TECTONIC VARIABILITY ALONG THE MAR

32o18'O"W

32 ° 16'30"W

32" 15'O'~V

32 ° 13'30"W

37 o20'0"N

37 ° 20'O"N

37 ° 19'O"N

37`, 19'0"N

37 ° 18'O"N

37 ° 18'0"N 32 ° 18'0WM

32 ° 16'30"W

32" 15'O"W

32" 13'30Wq

a) segment centre 32`,2 l'0'"OV

32 ° 19'30"W

32018'0"W

37o9,0,,N

37o9,0,,N

37 °8'O''N

37°8'0"N

w

32 ° 2 I'O"W

32 ° 19'30"W

32" 18'O"W

b) segment end (south) Fig. 5. Detail of TOBI images from (a) the centre and (b) southern end of Lucky Strike (PO1) segment. Bright tones correspond to strong backscattering surfaces, such as fault scarps and fresh volcanic outcrops. Darker tones correspond to low backscattering surfaces, such as low rugosity areas and pelagic sediments.

ubiquitous sedimentary cover. Sparsely distributed haystacks ( < 2 0 m high) and pillow lava flows overly the more fluid, sheeted lavas. Cracks, open fissures, fault scarps and fresh talus ramps are commonly observed at this part of the segment (Fig. 4b[I).

The southern part of the inner valley was mapped during dive OT 11. Here the axial valley is almost 8 km wide with a rough axial depth profile rapidly decreasing towards the nodal basin (Figs 2c, 4bl). The area is dominated by very fresh, coalesced pillow lava mounds and

8

E. GRJIC|A E T A L . 32 t8W

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1

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3718N

|

|

32 15W

32 13'30W

----

a) 32 21W

32 t9'30W

32 18W

32 21W

32 19'30W

32 18W

3709N

37 08N

37 07N

b) Fig. 6. Geological interpretation of Fig. 5. (a) The central platform is composed of sheet-flow lavas with a closely spaced faulting and lineated fabric. Volcanic breccia, pillow lava mounds and talus scree have been identified by correlating with Nautile observations at the central part of Lucky Strike (DIVA cruise, Fouquet et al. 1994; Ondr~as et al. 1997). (b) Towards the southern segment end, constructional volcanic features with characteristic hummocky textures (e,g. Smith et al. 1995) and localised deformation are observed.

volcanic ridges, with tectonic structures far less prominent than in the northern and central parts of the segment. Conical volcanoes, ranging from 2 0 m to 2 8 0 m high, are formed by large lava

tubes at the flanks and smaller haystacks and pillows at the summits (Fig. 4blI). Sedimented sheet-flows are locally uplifted, fractured and tilted forming shallow domes up to 20 m in

VOLCANO-TECTONIC VARIABILITY ALONG THE MAR diameter. These structures are often filled and covered by fresh pillow lavas suggesting recent local eruptions. Two Nautile dives (OTI2 and OT13) explored the rift valley walls (around 33°55'N) of the segment centre (Figs 2c, 3b). While the Eastern Wall is dominated by three main spaced steps (up to 450 m high each), some of them formed by large-throw single faults, the Western Wall is characterized by a succession of closely spaced, small-throw (10-50m high) fault scarps (Fig. 3b). The fault surfaces expose truncated pillow lavas, striated basaltic breccias and thick sections of massive basalts. At the foot of the lower fault scarps, there is fresh basaltic rubble suggesting recent tectonic activity. Gently outward dipping and heavily sedimented terraces are observed between the faults. The terraces on the Eastern Wall are several hundreds of metres wide and present a horst and graben-like microtopography as well as dispersed pillow lava mounds. L u c k y S t r i k e segment (PO1)." high resolution sidescan sonar ( T O B I ) results

Two TOBI sidescan sonar swaths covering the eastern rift valley wall and half of the axial valley floor were acquired along the Lucky Strike segment (Fig. 2a). Most of the central platform surface is characterized by a mottled/ patchy backscatter pattern (Fig. 5a). This was sampled by dredging on the LUSTRE 96 cruise (Fornari et al. 1996) and a series of glassy basalts and sheet-flow lavas were recovered (Grier pers. comm.). However, some of the fi-eshest basalts reported from anywhere in the whole segment are located in the collapsed lava lake central to the seamount complex, overlying the sheet-flow platform (Fouquet et al. 1995). This outcrop, however, is very restricted in volume and extent (300m diameter). Beyond the margins of the platform to the north and the south, the TOBI data show extensive linear clusters of constructional volcanic ridges oriented at about 018 °020 ° consisting of pillow mounds. This mounded terrain is locally cut by fault arrays, but the style of faulting changes systematically with increased distance from the platform. The faults which occur closest to the ends of the segment have consistently greater apparent throw and are more widely spaced than those outcropping in the middle of the segment. Centrally the faulting is more closely spaced, in some cases at less than 50 m intervals, and imparts a pervasive, intensely lineated fabric to the axial floor (Figs 5a, 6a). Towards both ends of the segment, beyond the

9

constructional volcanic features, patches of uniformly brightly-backscattering seafloor occur with irregular margins, suggesting that unfaulted sheet flows overlap the sedimented areas which dominate the NTDs (Figs 5b, 6b).

Discussion Crustal accretion processes and magma supply can be inferred from both dive observations and TOBI data, the latter providing synoptic coverage. Segment OH1 is characterized by centrally located extensive volcanic activity with outcrops of fresh sheet-flow lavas. Away from this central part little, if any, extrusive activity was detected during the dives acquiring data from the axis, and faulting and the extent of sedimentary cover increase significantly. The broad negative MBA 'bull's eye' situated at the segment centre (Derrick et al. 1995) suggests a centrally located focused mantle diapiric source (Lin et al. 1990). This correlates broadly with the variation in seismic crustal structure along the segment (Derrick et al. 1997), which reveals a higher velocity centrally than at similar depths at the segments ends. This variation is the result of enhanced intrusive activity at the segment centre in contrast with greater faulting near the segment ends (Fig. 7). In contrast, the shallowest and narrowest portion of segment OH3 is not located at the midpoint but is displaced 10kin towards the north. Far from corresponding to the region with present-day maximum volcanic activity, it is at this narrow section of the axial zone where tectonic activity predominates. This area is also where the thickest crust is imaged by gravity data (GrAcia et al. 1997a). Minor unsedimented volcanic extrusives interpreted as relatively recent are restricted to the southern section of the segment in a region of greater depth and thinner crust. We interpret this to suggest either a recent along-strike shift in locus of magmatic source, or a representation of the last vestiges of volcanism fed from lateral conduits within the shallow crust of the segment (Grficia et al. 1997a) (Fig. 7). The apparently diverse and complex pattern of volcanic extrusives observed along Lucky Strike has some along-axis symmetry which may provide clues to the shallow crustal magmatic supply system. Nonetheless, this symmetry of extrusive types could be interpreted in a number of different ways. One possibility is that a shallow crustal magma body has sourced the central platform flow series in a continuous fashion, while supplying (either through lateral dyke injection or marginal conduits) the non-

~eqment

"~eg e

Segment

~egment

Lucky Strike (P01)

Segment centre

Segment OH3 Segment end

Fig. 7. Interpretative sketch showing the along-axis volcano tectonic wmability of segment OH I, Lucky Strike (PO 1) and OH3. In the three cases, a single focused magma source is being tapped at a time. Surface geology is based from synthesis of the Nat;tile and TOBI data. Sections at depth are based on crustal thickness variations inferred From gravity data (Detrick eta/. 1995) and modified from the model of slow-spreading ridge segment by Cannat eta/. (1995). See text for further explanation.

Segment

Segment OH1

"n

"3

VOLCANO-TECTONIC VARIABILITY ALONG THE MAR 33° t 7'O"W

33 ° 16'O'~

33 ° 15'0'~

33 ° 14'0'~, ~

36°47,0,%

36~47,0.N

36~46'0"N

36*46tO,.N

36o45'0-N

36~45'0"N

33o l~,o,:,w

3~oi ~,o,,w

33o~'o"w

33° t ~,0~

a) segment centre 33 ~ t 5'O%~

33 ° t4~0'~'

33 ~ I 3 ' 0 " W

33 ° t 2'0'3N

33 ~ t I~OCW

36°55'0"N

36o55,0,,N

36 ° 54'0"N

36 ~54'0"N

36%3'0"N

36 ~53'0"N 33 ° 15'O"W

33 ° 14'0'~¢~t

33 ~ t 3'0%~/

33 ~ 12'O~W

33 ~ 1 I'O"W

b) segment end (north)

Fig. 8. Detail of TOBI images from (a) segment centre and (b) northern end of the FAMOUS (PO3) segment. Same key as Fig. 6.

central, more localized constructional volcanism and more distal sheet-flows (Figs 5b, 6b). Another possibility is that there are shallow magma sources throughout the segment, and the along-segment variation of extrusive type is a result of different thicknesses of crust providing varied lithostatic head, and hence varied extrusive forms. We do not have access to

petrological analyses which would help in differentiating between these scenarios, although recent sampling programmes could be useful in this respect (Langmuir, Grier, pets. comm.). Following extensive sidescan surveying and bottom observation programmes (Fornari et al. 1996) the recent volcanism has been reported only from the central lava lake area and one

12

E. GR,~CIA E T AL. 33 I?W

33 16W

33 15W

33 14W

36 47N

3647N

3646N

3646N

36 45N

3645N ! 3317W

I 33 16W

I 3315W

I 33 14W

a) 33 t5W

33 14W

3313W

3312W

331 lW

33 t5W

3314W

3313W

33 t2W

33 11W

36 55N

36 54N

36 53N

Fig. 9. Geological interpretation of Fig. 8. A continuous axial volcanic ridge (AVR) can be observed all along the segment (a and b). (a) The Uranus and Pluto mounds, explored during the FAMOUS project, are part of this continuous AVR. They are composed of coalesced fresh pillow lava edifices (Ballard & Moore 1977).

interpretation of restricted extrusive activity is that the segment is currently at a period of suppressed, perhaps declining magmatic activity. The persistently axially focused tectonism of the platform appears to further suggest that the segment is entering a tectonically-dominated phase. Further work (Fornari, Humphris, re-

search in progress) on the detailed relationships between neovolcanics and neotectonic phases will help determine whether the current balance is part of an evolutionary process or steady state. The symmetry of the extrusive volcanic pattern described for Lucky Strike has some

VOLCANO-TECTONIC VARIABILITY ALONG THE MAR direct and indirect commonality with each of the OH1 and OH3 segments (Fig. 7). The relative flatness, uniform acoustic texture, and lateral extent of the platform area central to the Lucky Strike segment suggest a construction from sheet-flow events. Its large volume further indicates a localized conduit (or set of conduits) as a source persistent over some time. Outside of this, a series of constructional volcanics and sheet-flow events occur up to 20kin from the centre of the segment. On TOBI records, these features have high backscatter levels and only localized deformation, suggesting that they are relatively recent events, as the events found in southern OH3. The sheet-flows which have been identified as comprising the platform area have a varying amount of sediment cover, largely dependent on their location with respect to prevailing bottom currents. Our interpretation is, therefore, that the lava lake represents a late stage extrusive event, post-dating the cessation of the more extensive constructional and sheetflow basalts. In this respect, Lucky Strike represents an intermediate position between OH I and OH3 (Fig. 7). In each of these segments, there is insufficient off-axis bathymetric data to determine whether the three segments studied are part of a time series, and as such representing 'snapshots' of an evolutionary cycle, or whether each of the segments is at a steady state of maintained magmatic level. We recognize that the interpretation presented here of the shallow crustal magmatic supply system (Fig. 7) is a simplistic one, and relies on a number of assumptions. The first being that at depth, a single, persistent source is tapped. The geometry of the source may differ from this assumed form, such as may be in the case of the F A M O U S segment (PO3), located just 25 km south of Lucky Strike (Fig. 1). The F A M O U S segment is 47 kin long with an 'hour-glass' shaped rift valley (2 to 6 km wide), and is characterized by a mostly continuous, shallow and fresh axial volcanic ridge (AVR) throughout the length of the segment (Figs 8, 9). In detail, this continuous AVR located towards the western valley flank, is formed by coalesced volcanic edifices up to several hundred metres high and up to 4.5 km long (Fig. 8a). These edifices correspond to the mounds (Venus, Pluto, Uranus and Mars) explored during the FAMOUS and A M A R (Crane & Ballard 1981) expeditions, which are composed of fresh pillow lavas, lava tubes and lobate lavas (Ballard & Moore 1977). In the gravity data, a weak ( - 5 m G a l ) and linear negative MBA extends throughout much of the length of the segment (Derrick et al. 1995). In this case, the pattern of

13

extrusives sourced at a number of points along the length of a more elongate magma body (Stakes et al. 1984) appears to correlate with our observations of a continuous, freshly constructed AVR to the western side of the FAMOUS segment axial valley floor.

Conclusions While there are obvious difficulties in extrapolation from photoreconnaisance and ground truthing of TOBI backscatter data, the scales of these methods of data acquisition make them suitable for increasing our understanding of the relationship between volcanic construction and tectonic activity, given adequate coverage. In this geographical location the extensive research - published and continuing - further aids interpretation and the development of models of crustal accretion. (1) Three contrasting second-order ridge segments of the Mid-Atlantic Ridge between 37°30'N and 33°30'N have been studied. Two of the segments were explored with the deep submersible Nautile (segments OHI and OH3) and the third (Lucky Strike segment, PO1) with the high resolution side-scan sonar TOBI. (2) The three segments appear to illustrate different stages in magmatic-tectonic processes of segment evolution. Segment OH1 can be considered as the classic 'model' of slow-spreading ridge segment, with the most recent volcanism and focused accretion located at the centre of the segment. In contrast with segment OH i, segment OH3 is characterized by a non-central magmatic source. Both shallowest bathymetry and maximum crustal thickness are skewed towards the north whereas neovolcanism is restricted towards the south of the segment. We interpret this as a recent re-location of the magmatic source or representing the last vestiges of a northern source fed by lateral injection. Lucky Strike segment represents an intermediate position between segments OH1 and OH3, with only localized recent magmatism at the segment centre, indicating the decline of a centrally located magma source. (3) In the three examples, we consider a single magma source being tapped at a time (or if multiple sources exist, they are tapped diachronously). This geometry may differ from other segments, such as FAMOUS, where a continuous magma source or several individual melt batches ascending

E. GRACIA ET A L.

14

b e n e a t h the axis a p p e a r to correlate with our seafloor observations. Support for the OCEANAUT cruise came from the D{partement des G{osciences Marines of IFREMER and the D~partement des Sciences de la Terre , UBO (Universit~ de Bretagne Occidentale) under the sponsorship of the research group GDR-GEDO. The cruise was part of the French/US FARA and InterRidge programmes, Sidescan sonar and interpretation was jointly funded by the UK-'BRIDGE' Programme and by EC-MAST2 'MARFLUX/ATJ' contract (MAS2 CT93-0070) co-ordinated by H, Bougault (IFREMER). E. Grficia benefited from a EC-TMR Marie Curie fellowship (project ERBFMB1CT950345). We thank H. D. Needham for the FARA-SIGMA cruise data, basis for the OCEANAUT and MARFLUX projects. The captain, officers and crew of the R/V Nadir as well as the Nautile Team are thanked for making this study possible. The scientists, participants of the OCEANAUT cruise (R. Hekinian, C. Bollinger, M. Constantin, C. Guivel, and B. Sichler) are thanked for their contributions in collecting the dive data. We are grateful to Y. Lagabrielle and R. Pallas for helpful discussions, and to particularly constructive critical reviews of two anonymous referees.

References

ARCYANA 1975. Transform fault and rift valley from bathyscape and diving saucer. Science, 190, 108-116. BALLARD, R. D. & MOORE, J. G. 1977. Photographic Atlas of the Mid-Atlantic Ridge Ri£t Valley. Springer Verlag, New York,. BALLARD, R. D., BRYAN, W. B., HEIRTZLER, J. R., KELLER, G., MOORE, J. G., & VAN ANDELT. 1975. Manned submersible observations in the FAMOUS area: Mid-Atlantic Ridge. Science, 190, 103-108. BIDEAU, D., HEKINIAN, R., BOLLINGER, C., ET AL. 1996a. Submersible observations of highly contrasted magmatic activities recorded along two segments of the Mid-Atlantic Ridge near 34°52'N and 33°55'N. Inter-Ridge News, 5(1), 9 14. -, -, ET' +4L. 1996b. Contrasted volcano-tectonic activities at 34°N and 35°N on the Mid-Atlantic Ridge: A submersible study. Journal qf ConJerence Abstracts, 1(2), 756-757. CANNAT, M., MEVEL, C., MAIA, M. E T A L . 1995. Thin crust, ultramafic exposures, and rugged faulting patterns at the Mid-Atlantic Ridge (22°N-24°N). Geology, 23. 49 52. CHARLOU, J. L., BOUGAULT,H., DONVAL, J. P., PELLE, H., LANGMUIR, C. ~:; FAZAR SCIENTIFIC PARTY 1993. Seawater CH4 concentrations over the Mid-Atlantic Ridge from the Hayes F.Z. to the Azores triple junction. EOS Transactions O/" the American Geoph)'sical Union, 74, 380. CRANE, K. ~¢ BALLARD, R. D. 1981. Volcanics and structures of the Famous narrowgate rift: Evidence for cyclic evolution: AMAR 1. Journal of Geophysical Research, 86, 5112 5124.

DEMETS, C., GORDON, R. G., ARGUS, D. F. & STEIN, S. 1990. Current plate motions. Geophysical Journal of the Royal Astronomical Society o[ London, 101, 425-478. DETRICK, R. S., NEEDHAM, H. D. & RENARD, V. 1995. Gravity anomalies and crustal thickness variations along the Mid-Atlantic Ridge between 33°N and 40°N. Journal of Geophysical Research, 100, 3767 3787, , COLLINS, J., KENT, G. Er AL. 1997. MidAtlantic Ridge Bull's Eye experiment: A seismic investigation of segment-scale crustal heterogeneity at a slow spreading ridge. Inter-Ridge News, 6(1), 27 32. FLEWELLEN, C. G., M1LLARD, N. W. & ROUSE, I. P. 1993. TOBI, a vehicle for deep ocean survey. Electronic and Communications Engineering Journal, 5, 85-93. FORNARI, D., HUMPHRIS, S. E. & SCIENTIFIC PARTY 1996. LUSTRE'96 cruise report. Multidisciplinary investigations of hydrothermal vents on Luclg,, Strike seamount attd the tectonic and volcanic structure of the Mid-Atlantic RMge Rift Vall
VOLCANO-TECTONIC VARIABILITY ALONG THE M A R LANGMUIR, C. H., DESONIE, D., REYNOLDS, J. ET AL. 1993a. Intensive rock sampling of 16 ridge segments between 32 ° and 41°N: preliminary results from shipboard analyses. EOS Transactions of the American Geophysical Union, 74, 380. , FORNARI, D., COLODNER, D. E/~ AL. 1993b. Geological Setting and characteristics of the Lucky Strike vent field at 37°17N on the MidAtlantic Ridge. EOS Transactions of the American Geophysical Union, 74, 99. LIN, J., PURDY, G. M., SCHOUTEN,H., SEMP~RE,J. C. 8¢ ZERVAS, C. 1990. Evidence from gravity data for focused magmatic accretion along the MidAtlantic Ridge. Nature, 344, 627 632, NEEDHAM, H. D., VOISSET,M., RENARD,V., BOUGAULI, H., DAUTEIL, O., DETRtCK, R. & LANGMUIR, C. 1992. Structural and volcanic features of the MidAtlantic rift zone between 40°N and 33°N. EOS Transactions o[" the American Geophysical Union, 73, 552. OONDRI~AS,H., FOUQUET,Y., VOISSE'i-,M. & RADFORDKNOERY, J. 1997. Detailed study of three contiguous segments of the Mid-Atlantic Ridge,

15

South of the Azores (37°N to 38°30N) using acoustic imaging coupled with submersible observations. Marine Geophysical Researches, 19, 231-255 SEMPERE, J. C., LIN, J., BROWN,H. S., SCHOUTEN,H. & PURDY, G. M. 1993. Segmentation and morphotectonic variations along a slow spreading center: The Mid-Atlantic Ridge (24°N-30°40'N). Marine Geophysical Researches, 15, 153-200. SMITH, D. K., CANN, J. R., DOUGHERTY,M. E. ET AL. 1995. Mid-Atlantic Ridge volcanism from deeptowed side-scan sonar images, 25°-29°N. Journal of" Volcanology and Geothermal Research, 67, 233262. STAKES, D. S., SHERVAIS,J. W. ~; HoPsoN, C. A. 1984. The volcano-tectonic cycle of the FAMOUS and AMAR valleys, Mid-Atlantic Ridge (36°47'N): Evidence from basalt glass and phenocryst compositional variations for a steady state magma chamber beneath the valley midsections, A M A R 3. Journal Geophysical Research, 89, 6995 7028.

A seismic inversion study of the axial magma chamber reflector beneath the East Pacific Rise near 10°N J. S. C O L L I E R

& S. C. S I N G H

D e p a r t m e n t o f Earth Sciences, Bullard Labs, M a d i n g l e y Road, University o j Cambridge, Cambridge CB3 0 E Z , U K Abstract: We have applied a full waveform inversion method to seismic reflection data from

two adjacent fourth-order morphological segments of the East Pacific Rise near 10°N. Our inversion results show the ridge to be underlain by a thin (30 m) low P-wave velocity layer which we interpret as magma with less than 20+10% crystals. At one of the locations, the base of the mostly molten layer is underlain by a velocity gradient. Here the ridge is at an intermediate stage in its volcanic cycle and we interpret the gradient zone to be a result of an increase in crystals with depth, reaching 40 90% crystallinity over a vertical distance of 50 m. In contrast, beneath the second segment, which is at the onset of a vigorous volcanic stage, the mostly molten layer is underlain by a mostly solid floor. We infer that the melt layer has been newly emplaced. Our results provide support for the idea that the crustal magma system beneath fast spreading ridge axes is segmented on a similar scale to the fourth-order bathymetric segmentation.

A number of different indicators have been used to infer the tectonic and magmatic state of fast spreading mid-ocean ridges. These indicators include axial depth (Macdonald & Fox 1988); the presence or absence of an axial summit caldera (Macdonald & Fox t988); the crosssectional area and shape of the ridge crest (Macdonald et al. 1984); lava age (Haymon et al. 1991a); density and width of fissures (Wright et al. 1995), intensity of hydrothermal venting (Haymon et al. 1991a) and basalt MgO content (Langmuir et al. 1986). Taken together these studies have proposed a three stage cycle of volcano evolution. The beginning of the cycle marks the most active stage in which there is vigorous magmatic and hydrothermal activity. If a summit caldera were present before the onset of this phase it may become filled with fresh lavas and disappear. The middle of the cycle is characterized by continued hydrothermal activity at older, established sites but no seafloor eruptions. Wide fissures associated with dyke injection may develop. A summit caldera, thought to be formed by magma chamber roof collapse following the evacuation of melt in the earlier stage (Macdonald & Fox 1988), would be expected to be present. The final stage is marked by the cessation of magmatic and hydrothermal activity and the development of fine fissures due to cooling and tectonic extension. This cycle has been inferred to be driven by periodic delivery of magma from the upper mantle to particular sections of the axis (Macdonald & Fox 1988). Macdonald & Fox (1988) suggest that there is a correlation between parts of the ridge at the

end of a volcanic cycle and the absence of an ~Axial Magma Chamber' (AMC) reflector. The A M C reflector has been imaged at a depth of 1.5-3.0 km beneath a number of different intermediate to fast spreading ridges (Detrick et al. 1987; Morton et al. t987; Collier & Sinha 1990; Mutter et al. 1995). The current interpretation of this event is that it is the reflection from a thin layer of mostly molten material which overlies a broader region (~6 kmwide) of 'hot rock' (Pwave velocity anomaly of up to 1 km s-1) with at most ~ 3 - 5 % partial melt fraction in the lower crust (Sinton & Detrick 1992). Note that throughout this paper we use the term ' A M C ' to refer to this thin layer of mostly molten material only. It has been suggested that solidification and subsidence of the floor of the thin melt layer forms the cumulate sequence of the lower crust (Sleep 1975: Browning 1984; Phipps Morgan & Chen 1993; Henstock et al. 1993). If this is correct then the A M C fulfils a central role, not only as a reservoir of material to form the upper crust, but also in the accretion of oceanic crust as a whole. To date, the available seismic reflection data, although extensively used for mapping purposes, has only been analysed to determine the detailed structure of the melt layer in a few isolated locations (Kent et al. 1993; Phipps Morgan et al. 1994; Hussenoeder et al. 1996). In addition, the previous studies have just used forward modelling to match the m a x i m u m peak-to-peak amplitude with offset and approximate form of the reflector with the fixed assumption that the A M C is caused by a simple threedayer sill (solid

COLLIER, J. S. & S~Nct~, S. C. 1998. A seismic inversion study of the axial magma chamber reflector beneath the East Pacific Rise near 10°N. In." MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 17-28

17

18

J. S. COLLIER & S. C. SINGH

Bathymetry(m) . 2 5 0 0 2 6 0 0 2700 2 8 0 0 2900 3 0 0 0 3100 3 2 0 0 3300 3400 3500

9" 54'N

9 ° 48'N

9 ° 42'N

9 ° 36'N

9 ° 30'N

9 ° 24'N

9 ° 18'N

104 ° 30'W

104 ° 24'W

104 ° 18'W

104" 12'W

104 ° 06'W

9 ° 12'N 104 ° 00'W

Fig. 1. Bathymetric map of the East Pacific Rise showing locations of seismic data. Across-axis reflection profiles (labelled with line number and width of the AMC reflector from Kent et al. 1993) are shown with fine lines and the along axis seismic reflection profile with a bold line. Stars mark the locations of the CDP data used in this study. The location of the mid-point of ESP 5 used for the long-wavelength velocity profile is shown with the black square. The dashed line box marks the location of the seismic tomography experiment (Toomey et aI. 1990). Devals are shown with white triangles. The full spreading rate at this location is 11 cm a 1

roof, melt layer, solid floor). This model was devised by the lack of a second distinct reflector below the A M C reflector (or 'basal' reflector) which could be due to the sill being less than a seismic w a v e l e n g t h thick ( 1 5 0 - 3 0 0 m for a typical experiment), such that the reflection from its base interferes with the reflection from its top. H o w e v e r , K e n t e t al. (1993) also acknowledge that the lack of a basal reflector could be due to a gradient zone beneath the molten layer. Collier & Singh (1997) present the first application of a waveform inversion m e t h o d to seismic reflection data collected at the East Pacific Rise. In this application we free ourselves

of any pre-set, biased solution by using an inverse m e t h o d to solve for structure directly from the seismic data themselves. Our m e t h o d is an a u t o m a t e d one, enabling us to search through literally hundreds of models to determine that which best fits the complete wavefield i.e. all amplitude and phase characteristics of every frequency contained in the data. This paper extends the w o r k previously p r e s e n t e d by comparing results from different sections of the ridge axis system inferred from ocean b o t t o m observations to be at different stages in their magmatic cycle and hence assessing the influence of the A M C on volcanic processes.

SEISMIC STUDY OF THE EPR MAGMA CHAMBER

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Fig. 2. Above: Fourth-order Tectono-magmatic segments and interpreted stages of volcanic cycles according to Haymon et al. (1991a) and Wright et al. (1996). Below: Stacked multichanne! seismic reflection profile collected along axis (Detrick et al. 1987). See Fig. 1 for line location. The vertical solid lines show locations of CDP gathers used in this analysis. The vertical dashed lines show locations of bathymetric devals. The box shows the section shown in Fig. 3. Positive polarities are shaded black.

Study area We selected the area of the East Pacific Rise between 9°12'N and 9°54'N (Figs 1 & 2). This part of the ridge axis has been intensively studied, with the collection of multibeam bathymetry (Macdonald et al. 1984), side-scan sonar (SeaMARC II, Macdonald & Fox 1988; A R G O Haymon et al. 199 la), video camera (Haymon et al. 1991a), submersible dives (Haymon et al. 1991 b) and systematic sampling (Langmuir et al. 1986; Batiza & Nui 1992). On the basis of these bottom-observation studies a number of different, ~ 1 0 k m long ridge sections--known as 4th order segments--each in a distinct stage of the established volcanic cycle (~1000a) have been recognized (Haymon et al. 1991a). Some of these morphologic segments are bounded by small deviations of axial linearity (devals) of the axis and others by changes in the axial summit caldera. Conventional single ship seismic reflection and two ship expanding spread (ESP) profiles were collected along this section of ridge in 1985 (Fig. 1; Detrick et al. 1987; Kent et al. 1990, 1993; Vera et al. 1990). We reproduce the alongaxis seismic reflection profile together with the recognized morphologic segments (numbered A - G ) in Fig. 2. The A M C reflector is seen between 4.1 and 4.3 s TWT (~0.6 s TWT below

the seafloor) along much of the profile. The across axis seismic reflection profiles showed the width of the reflector to be just 0.25M.15km (Kent et al. 1993). Devals appear to coincide with distinct jumps in either the width of the reflector or its position relative to the ridge crest. However, the fourth-order discontinuities themselves were not consistently associated with narrow or absent AMCs as envisaged by some models (e.g. Macdonald & Fox 1988) of melt distribution beneath axes. Comparison with images obtained on across-axis lines showed that the disappearance of the reflector from the along-axis image (known as dropouts) is associated with navigational errors--the narrowness of the reflector making it an easy target to miss. Kent et al. (1993) concluded that the reflector underlay the whole of the 80kin ridge portion surveyed, but varied in width and/or lateral position beneath various morphological segments. Seismic tomography between 9o26 '9°35'N (Segments D-F) also showed the morphological segmentation (devals at 9°28'N and 9°35'N) to correspond to segmentations in crustal velocity structure, leading Toomey et al. (1990) to suggest that small-scale morphologic segmentation of the East Pacific Rise is magmatically controlled. Our analysis assumes a one-dimensional structure so we needed locations where the

20

J. S. COLLIER & S. C. SINGH

3.4

3.6

v

I--

3.8

4.0

4.2

lO2OO

lo4oo

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losoo

11000

Fig. 3. Enlargement of along axis-stacked seismic reflection profile shown in Fig. 2 but plotted with reverse polarity to emphasize along-axis variations in the waveform of the AMC. Negative polarities are shaded black. Data were averaged between the two vertical lines at each chosen CDP site.

A M C was centrally crossed on the along-axis line and where the waveform of the A M C was consistent over a distance of at least a few km and where the seafloor is locally flat. This ruled out the southern part of the line where the ship track did not follow the crest of the ridge well (Fig. 1). Also the central part of the line (near 9°30'N) was considered unsuitable, as although the A M C is centrally located beneath the topographic axis the underlying low-velocity zone is offset about 1 km to the west (Toomey et al. 1990; Kent et al. 1993). We therefore confine our investigation to magmatic segments B and C. Segment B is interpreted as being at the beginning of a volcanic cycle. It is the highest point between Clipperton and the large 9°03'N overlapping spreading centre that forms the southern boundary of this second-order ridge segment. The average age of the lavas is less than 50 a. The section between 9°46'-9°51'N was observed erupting pillow basalts by submersible dives in 1991 (Haymon et al. 1991b). Segment B2 lacks a summit graben (Macdonald & Fox 1988; Haymon et a/. 1991a) and is the least fissured and most hydrothermally vigorous (both white and black smokers) stretch of the ridge between 9°12 ' and 9°54'N (Wright et al. 1995). In contrast, Segment C is thought to be in the middle of a cycle. It is characterized by wide (5-20 m) fissures, a wide (150-210m) and deep (15m) axial summit caldera and young, but not fresh, basalts thought to be 50-100 yr old.

Analysis We show an enlargement of the stacked seismic section collected over segments B and C in Fig. 3. In this figure we have plotted the data with reverse polarity to emphasize the along-axis profile variations of the waveform of the AMC. We selected two areas for analysis, C D P (Common Depth Point) 10340 (9°39'N) and C D P 11050 (9°48'N), where the waveform was consistent over a lateral distance of at least 3 km. The stacked section clearly shows the A M C waveform to be different at these two chosen locations---consisting of a single black stripe around C D P 10340 and a double black stripe around CDP 11050. CDP 11050 is characteristic of the A M C underlying the northern part of segment B2 and segment B1. CDP 10340 is characteristic of the whole of segment C and the southern part of segment B2. To improve signal-to-noise and also to minimize small two-dimensional effects due to variable upper crustal structure we made a 55fold constant-offset stack in each of the two selected areas to produce super-gathers. This is equivalent to averaging over a lateral distance of 1.4km. We then transformed the gathers into the intercept time-slowness (tau-p) domain for ease of interpretation. The results are shown in Fig. 4--the difference in the A M C waveforms between the locations is again clear. At C D P 11050 the reflector is a triplet, with three large peaks such that it looks like a 'W'. At C D P

SEISMIC STUDY OF THE EPR MAGMA CHAMBER

21

CDP 11050

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Fig. 4. Super-gathers (55-fold constant offset stacked) showing the waveform of the AMC reflection at our two chosen sites. The data are shown in the intercept time-slowness (tau-slowness) domain. Positive polarities are shaded black. To the left are the source wavelets used in the inversion at each location.

10340 the reflector is a doublet, with two large peaks. A critical question in our analysis is whether the difference in the waveforms at the two locations is real or a seismic imaging-related artefact. Seismic data that are not collected across the strike of the indented target can suffer from sideswipe, i.e. contamination from out-ofthe-plane events. We therefore have the possibility that drifting of the 2.4km hydrophone streamer with respect to the A M C target is responsible for the variation in waveform we observe. With the available, sparse 2D arrangement of seismic profiles we cannot unequivocally rule this possibility out. However, we have two lines of evidence, in addition to the mapping results of Kent et al. (1993) and along axis consistency of waveform which suggest this is not the case. Firstly we compared the AMC arrival times at our locations with those on the closest unmigrated across-axis section to ensure that they were from the central part, and not the diffractive tail, of the reflector. Secondly, Kent et al. (1993) describe a relationship between AMC

reflection strength and ship location (the amplitude falling where the target was missed)--both our gathers have a similar, high amplitude. It was also important to rule out the possibility that the difference in the waveforms was due to a different source signature. The airgun array used in this experiment consisted of just four guns, so any small variation in towing depth or timing of one of these guns had the potential to change the signature significantly. We therefore extracted the source wavelet for each location independently (see Collier & Singh 1997 for details) to ensure this was not a problem (Fig. 4). The application and limitations of our inversion method are fully described in Collier & Singh (1997) and summarized in Fig. 5. As input to the inversion we provide an estimate of the source wavelet, an initial velocity model and the slowness tau transformed data. The source wavelets were each 250 ms in length so as to include the first two bubble pulses. The initial model defines P-wave velocity (Vp), S-wave velocity (Vs), density (p), P-wave attenuation (Qp) and S-wave attenuation (Qs) for a stack of

22

J.S. COLLIER & S. C. SINGH

I Initialvelocitymodel1 {S . . . . . . . . . let]

:~ F. . . . d model ~:

10%) compared to large changes in Vp (drops up to 60%) and Vs (drops up to 100%). We estimate that errors in our solutions are +8 m in depth and =t=0.2km s i in P-wave velocity.

( Syntheticseismog....

Results Seismicdata ~

Computemisfitfunction YES

Testfor convergence]

~NO ] Applyconjugategradientalgorithi~ Perturbvelocitymodelt

~-l-~

Outputvelocitymodel1

Fig. 5. Schematic flow chart of the seismic waveform inversion procedure applied in this study. We calculated synthetic seismograms for a horizontally layered model using the generalized reflection transmission matrix method (Kennett & Kerry 1979). This method accurately treats multi-pithing and mode conversions. We used a conjugate-gradient method given by Kormendi & Dietrich ( 1991 ).

8m thick isotropic, elastic layers. We used the model of Vera et al. (1990) obtained from modelling ESP data (Fig. 6a) for our initial model. We also performed inversions with the attenuation structure of Christeson et al. (1994) but this made negligible difference to the final results. Throughout our analysis we scaled the source wavelet such that the peak-to-peak amplitude of the seabed reflection matched that of the synthetic seismograms generated with the seafloor P-wave velocity set to 2.2 km s ~ (Vera el al. 1990). The fit of the synthetic seismograms generated with our initial model to the data is good (Fig. 6b) except for the waveform of the A M C itself. During the inversion we update Vp whilst keeping the other model parameters fixed. The data have a recording window of 0.32.7kin, which restricts our resolution of Vs structure below the top of the AMC. However, as the reflection amplitude-with-offset behaviour at the top of the melt body is strongly dependent on the difference in Vs across the interface, we can discriminate between simple models in which the melt is highly molten (Vs < 1.0 km s-1) and only partially molten (Vs> 1.0kms-1). It was unnecessary to attempt to invert for p as Murase & McBirney (1973) show that density changes little during melting (drops less than

Our final velocity models for each CDP location are shown in Fig. 7. Note that we show only the part of the model and seismograms for the A M C itself (the initial model gives a good fit to the upper crustal reflectors and so changed little as the inversion progressed). The fit of the synthetic seismograms to the data at C D P 11050 is slightly poorer than at CDP 10340. We attribute this to greater lateral variability in the stack region of C D P 11050 (Fig. 3). The fit at C D P 10340 is excellent. The inverted solutions for A M C structure have both similarities and differences at the two locations. Both showed the roof of the melt body to be a sharp velocity discontinuity. We tested models in which we introduced a more gentle gradient at the roof but these resulted in unsuccessful inversions. The roof of the chamber is about 100m shallower for the northern location (CDP 11050). Immediately below the roof, the melt layer has a Vp of 2 . 4 k m s 1 at CDP 11050 and 2 . 6 k m s q at C D P 10340. It is important to remember that seismic reflection is sensitive to acoustic impedance contrasts rather than absolute values. Therefore these absolute melt velocities have a dependency on the acoustic parameters of the roof region. The acoustic parameters of the roof are not well constrained in either this dataset or wide-angle data collected just to the south (Toomey et al. 1990; Vera et al. 1990). However the small Vp values obtained within the melt layer was supported by the low Vs velocity deduced within the melt body at both localities. If Vs of the melt was set above 1.0 k m s l (with that in the roof unchanged) inverting for Vp failed to find a model which satisfactorily generated the small amplitudes observed at farthest slownesses. Note that we are unable to resolve Vs structure below the top of the A M C and so do not show it below this in Fig. 7. Comparison with experimental results (see below) show that Vp in the range 2.4-2.6 km s-1 is consistent with near zero Vs for a basalt close to its liquidus. We therefore have some confidence that our absolute Vp values are reasonable. The low Vp layer has a thickness of about 30m at both locations. At the more southerly location (CDP 10340) the floor of the A M C has a velocity gradient, where Vp rises from 2 . 6 k m s -1 to 3.5kms -1 over a vertical distance of 50 m. We do not resolve Vp structure below this depth. In comparison the floor of the

SEISMIC STUDY OF THE EPR MAGMA CHAMBER

23

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AMC at the more northerly location (CDP 11050) has a sharp interface where Vp increases rapidly from 2.5 km s 1 to 5 km sq (similar to that of the roof). Again we do not resolve Vp below this depth.

Conversion o f seismic velocities to degree o f partial melt

Data on the variation of the seismic velocities of basalt and gabbro with degree of partial melting at pressures appropriate for mid-ocean ridges is limited. However, those studies that do exist (Christensen 1972; Murase & McBirney 1973; Khiratov et al. 1983; Manghnani et al. 1986) plus appropriately scaled (Henstock et al. 1993) analogous data on the high-pressure partial melting of peridotite (Sato et al. 1989) can be used to estimate the melt fraction from our seismic results. At first glance the most promis-

ing experiments to perform the conversion (in terms of sample composition and temperature range) are those of Murase & McBirney (1973) reproduced in Fig. 8a. However, these experiments were carried out under atmospheric pressure and experienced sub-solidus microcracking which reduced measured Vp below what would be expected, had confining pressure been applied. In addition, the measured velocities varied according to the cooling rate of the sample, indicating that cracking had a strong effect on the results, it is also unclear from these experiments whether textural equilibrium had occurred in the samples, so the melt configuration matched that which would be achieved in nature. Mavko (1980) demonstrates the importance of the configuration of the melt within the matrix on seismic velocity. The general form of the curve however is thought to be broadly correct (Phipps Morgan el al. 1994). We therefore chose to adopt the scheme of Henstock et

24

J.S. COLLIER & S. C. SINGH

Vel ( k m / s ) 0.0 2.5 5.0 'E4.0

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Fig. 8. Curves of Vp as a function of temperature. (a) Experimental results from Murase & McBirney (1973) for a basalt with 10% MgO. Vp at room temperature and pressure was 6.2 km s 1. (b) Normalized curves. The velocity is normalized to the velocity at the solidus temperate (~n), and temperature is normalized as t * = ( t - t m ) / ( t l - tin), where tm is the solidus temperature and tl the liquidus temperature, Solid line is the normalised curve of Henstock et al. 1993 for small degrees of partial melting. The dashed line is the data shown in (a) normalised with t m = 920°C; tl = 1200°C and vm = 6.2 km s ]. We used the solid curve with the dashed extension for higher degrees of partial melting to convert from Vp to t * and then used the curves of Sinton & Detrick (1992) to convert from t * to crystal fraction (see Fig. 9).

al. (1993). T o cover the full range o f velocities in o u r study we e x t e n d e d the n o r m a l i z e d curve o f H e n s t o c k et al. (1993, their Fig. 9b) (solid line Fig. 8b) by a p p r o p r i a t e l y scaling the experim e n t a l results o f M u r a s e & M c B i r n e y (1973) w i t h i n the m e l t i n g interval (dashed line Fig. 8b). T h e e x p e r i m e n t a l m e a s u r e m e n t s near the liquidus are less affected by the p r o b l e m o f confining

pressure a n d are therefore close to the truth. H a v i n g c o n v e r t e d Vp to n o r m a l i z e d t e m p e r a tures, we c o n v e r t e d n o r m a l i z e d t e m p e r a t u r e to melt fraction using the experimental data c o m p i l e d by S i n t o n & D e t r i c k (1992, their fig. 7b). We s h o w o u r results in Fig. 9 (hashed areas). As an alternative a p p r o a c h we also c o m p u t e d

SEISMIC STUDY OF THE EPR MAGMA CHAMBER CDP 10340

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80 60 40 20 0 % crystals

100 60 60 40 20 0 % crystals

Fig. 9. Two estimates of crystal content with depth inferred from our seismic inversion results. Hashed areas are bounds from experimental data of melting basalt and peridotite (computed from the curves of Fig. 8b). Dotted areas are bounds from a Hashin-Strickman calculation (see Collier & Singh 1997 for formulation) for a two phase material (melt: Vp=2.3kms 1, Vs=0.0kms-~, p=2.7Mg m 3; crystals: Vp--6.2kms 1, Vs=2.9kms-1 p = 2.8 Mg m-3) in two configurations (upper bound from unconnected melt inclusions in a solid host and lower bound from unconnected crystals in a molten host).

H a s h i n - S t r i c k m a n bounds for a two-phase medium (melt and crystals). The maximum bound comes from the case with unconnected melt inclusions in a solid host, the minimum bound from unconnected crystals in a molten host (see Collier & Singh 1997). Although these are unlikely geometries in nature they provide some bounds on the melt fraction indicated by our seismic velocity structures. We show our results in Fig. 9 (dotted areas). As we would expect the melt fraction with depth curves computed by our two approaches agree reasonably well at high and low melt fractions, but there is greater discrepancy in the partial melt region where melt topology forms an i m p o r t a n t control on seismic velocity (O'Connell & Budiansky 1977; Schmeling 1985). Although the errors in our conversion may be large (particularly in the middle-ranges of partial melting) we are confident that the broad pattern is correct, namely that a 30m thick mostly molten sill (crystals between 10 and 50%) is underlain by a more mushy, crystalline gradient zone (crystallinity rising to 40-90% over a vertical distance of 50 m) at CDP 10340, and a mostly solid base at CDP 11050.

Discussion M e l t layer thickness With no a priori assumptions we have shown the AMC to have the general form of a thin melt sill at the two locations at the East Pacific Rise near 10°N. We show the thickness of the sill to be of

the order of 30-80 m. Assuming the melt layer to be either positively (Stolper & Walker 1980) or neutrally (Ryan 1987) buoyant, and other factors such as pattern and degree of hydrothermal circulation being equal, we might expect from mechanical arguments the thickness of a melt layer to scale with its width. We note that the thinner (vertical thickness) A M C at 9°48'N (CDP 11050) is 200m narrower (across-axis width) than the A M C at 9°39'N (CDP 10340). An important question is how do our in situ determined melt layer thicknesses compare with estimates from field observations and theoretical studies. Arguably the best analogue for fast spreading crust is the Oman Ophiolite. Browning (1984) from a study of the cryptic layering of the cumulate section of the Oman Ophiolite proposed depth independent crystallization within discrete layers of melt about 100m thick. MacLeod (1997 pers. comm.) from structural and petrological mapping of the upper gabbro and dyke contact proposes the presence of a 200m thick mostly molten layer. Both these estimates are in approximate agreement with our in situ determination of melt thickness. According to theoretical models of melt compaction, magma is halted in its ascent when it reaches a 'freezing horizon' (the depth of the magma solidus), where the dilational volume change associated with magma freezing leads to viscous stresses that favour magma ponding within one viscous 'compaction length' (McKenzie 1985; Phipps Morgan & Chen 1993). For a picritic basalt separating from a 1% partial melt the compaction length is 100 m (McKenzie 1985)

26

J.S. COLLIER & S. C. SINGH

and hence compares favourably with our seismically determined thicknesses.

Melt volumes and implications to chamber connectivity The fact that one of the sections of ridge under study erupted just 6 a after the collection of the seismic reflection data presents the opportunity to perform a few simple calculations to estimate the size of the chamber that was tapped. Gregg et al. (1996) estimate the extruded volume of the 1991 eruption of segment B2 to be 5±1x106 m 3 and to have taken place over an 8 km stretch of ridge. If we assume that all the melt for this eruption came only from the underlying AMC (i.e. no along-axis lateral melt transport within the lens) and the AMC to be a rectangle, 500m wide (distance across axis), 8000m long (distance along axis) and 30 m deep and filled with eruptible material then the total volume of magma available was 1.2x 10Sm 3. If we assume that this melt volume was completely exhausted during the eruption and all dykes are vertical and reach the surface, the dyke volume emplaced would have been 10m wide (by 1400m high, by 8000m long). This is equivalent to 90 a of average plate separation. To achieve a more modest dyke injection width then we must assume that either the AMC was not exhausted in length, depth or breadth. For example, if we assume the total dyke width injected during this single eruptive episode to be 1 m, and the AMC was exhausted in depth and width then it would have been just 1000 m long. Similarly if it were exhausted in width and length then we would have tapped just 4m in depth. We conclude that either the eruption was halted by solidification within the chamber-to-surface conduit system or that melt within the chamber is either stratified or is not physically connected beneath the length and/or breadth of the ridge. We would be unable to detect either subtle density stratification as predicted by Sparks & Huppert (1984) or breaks in the AMC reflector smaller than 500m (the Fresnel radius). In addition, with the currently available data we are unable to ascertain whether the edges of the chamber are more solidified than the centres. This might well be expected on thermal grounds, and would limit the volume of melt that could be tapped during a given eruption.

Magma chamber processes Our results lead us to suggest there to be a relationship between the character of the AMC

and the tectono-volcanic stage of the ridge. Beneath segment B, which is currently erupting high temperature basalts at the seafloor, we have determined a rather simple, three-layer sill structure. The detection of a near-solid base to the sill suggests that significant cooling and crystallization had occurred (presumably primarily as the result of hydrothermal circulation at the end stage of the previous volcanic cycle) before the emplacement of a new batch of melt from the mantle. Beneath segment C, which is no longer erupting lava at the seafloor, we are observing the intermediate phase where the thin melt body slowly cools and freezes out. If we follow the arguments of Marsh (1989) who proposed that once crystallinity in a basaltic melt exceeds 25% it behaves rheologically like a solid, then the material from the gradient region is unlikely to be ejected to form part of the sheeted dyke sequence. The shape of the velocity-depth profile at 9°39'N (CDP 10340) implies that accretion of material to the roof of the sill is minor and that crystal settling under gravity is a far more important physical process. Our observations are at odds with the commonly held notion that vigorous cooling of the lens from above by hydrothermal circulation causes freezing or plating of gabbro from the lens onto the base of the sheeted dyke complex. We note, however, that limited crystallization at the roof of a body cooled from above is predicted by the theoretical models of Worster et al. (1990). Most of the samples analysed by Batiza & Nui (1992) contain about 10% phenocrysts, primarily plagioclase. They conclude that most of the lavas have experienced solid-liquid fractionation, with the addition of plagioclase (which they show will have a lower density than the basalt melt) and removal of mafic crystals (olivine and pyroxene). They suggest that gravitational settling is the most likely explanation for this process. They attempt to determine where the solid-liquid fractionation o c c u r r e d - either below the AMC, within the AMC, with the eruptive dykes, or after eruption. On thermal grounds they favour the AMC itself to the site of the fractionation. We interpret our velocity structure at 9°39N (CDP 10340) as a direct physical detection of such a process.

A segmented chamber and implications for magma supply geometry Previous seismic studies (Toomey et al. 1990; Kent et al. 1993) have suggested the seismic structure of this part of the East Pacific Rise to

SEISMIC STUDY OF THE EPR MAGMA CHAMBER be segmented on a scale similar to the 4th-order bathymetric segmentation, but there appears not to be a simple one-to-one geographical correspondence between the two. O u r results support these earlier studies by suggesting that adjacent 4th-order m o r p h o l o g i c a l segments are underlain by melt sills with different internal structure. A correlation between 4th-order bathymetric segments and: the width; location relative to the ridge axis; and reflection coefficient of the A M C was also established at the Valu F a Ridge in the L a u Basin (Wiedicke & Collier 1993). A key question is w h e t h e r a segmented crustal m a g m a c h a m b e r necessarily implies a segmented supply geometry. There is currently a heated debate as to w h e t h e r m a g m a supply f r o m the m a n t l e at fast spreading ridges is u n i f o r m alongaxis (2D) or confined to a few discrete centres (3D). If 3D, on w h a t length scale are the foci separated? M a c d o n a l d et al. (1988) p r o p o s e a 3D m o d e l in which injection foci are centred beneath f o u r t h - o r d e r b a t h y m e t r i c segments. Lin & P a r m e n t i e r (1988), however, argue that the small along-axis gravity signal (just 10 regal) and bathymetric variation at fast spreading ridges favours the 2D model or a 3D m o d e l only if the m a g m a c h a m b e r is temporarily persistent and well c o n n e c t e d along-axis. To reconcile the seismic results with those o f Lin & P a r m e n t i e r (1988) we suggest that either the supply is 3D on a f o u r t h - o r d e r segmentation scale but varies in space and time such that the integrated volume o f supply with time is uniform, or the m a n t l e supply is 2D but crustal variations relating to fracturing and h y d r o t h e r m a l fluid penetration, impose a 3D effect.

27

COLLIER,J. S. • SINHA,M. C. 1990. Seismic images of

a magma chamber beneath the Lau Basin backarc spreading centre. Nature, 346, 646-648. - & SrN~H, S. C. 1997. Detailed structure of the magma chamber beneath the East Pacific Rise at 9°40'N from waveform inversion of seismic reflection data. Journal of Geophysical Research, 102, 20287-20304. DETRICK, R. S., BUHL, P., VERA, E. E, MUSTER, J. C., MADSEN, J. A. & BROCHER, T. M. 1987. Multichannel seismic imaging of a crustal magma chamber along the East Pacific Rise. Nature, 326, 35-41. GREGG, T. K. P., EORNARI D. J., PERF1T, M. R., HAYMON R. M. & FINK, J. H. 1996. Rapid emplacement of a mid-ocean ridge lava flow on the East Pacific Rise at 9°46'-51'N. Earth and Planetary Science Letters, 144, E1-ET. HAYMON, R. M., FORNARI, D. J., EDWARDS,M. H., CARBOTTE, S., WRIGHT, D. & MACDONALD,K. C. 1991a. Volcanic eruption of the mid-ocean ridge along the East Pacific Rise crest at 9°45'-52'N: Direct submersible observations of the seafloor phenomena associated with an eruption in April 1991. Earth and Planetary Science Letters, 119, 85-101. , FORNARI, D., VON DAMM, K. ET AL. 1991b Eruption of the EPR crest at 9°45'-52'N since late 1989 and its effect on the hydrothermal venting: Results of the ADVENTURE program, an ODP site survey with Alvin. EOS Transactions of the American Geophysical Union, 72, 480. HENSTOCK, T. J., WOODS, A. W. & WHITE, R. S. 1993. The accretion of oceanic crust by episodic sill intrusion. Journal of Geophysical Research, 98, 4143~4161. HUSSENOEDER,S. A., COLLINS,J. A., KENT, G. M. ETAL. 1996. Seismic analysis of the axial magma chamber reflector along the southern East Pacific Rise from conventional reflection profiling. Journal of Geophysical Research, 101, 22 087-22105. KENNETT, B. L. N & KERRY,N. J. 1979. Seismic waves References in a stratified half-space. Geophysical Journal of the Royal Astronomical Society, 57, 557-583. KENT, G. M., HARDING,A. J. & ORCUTT, J. A. 1990. BATIZA, R. & NuI, Y. 1992. Petrology and magma Evidence for a smaller magma chamber beneath chamber processes at the East Pacific Rise the East Pacific Rise at 9°30'N. Nature, 344, 650,~9°30'N. Journal of Geophysical Research, 9% 653. 6779-6797. , , & 1993. Distribution of BROWNINO, P. 1984. Cryptic variation within the magma beneath the East Pacific Rise between cumulate sequence of the Oman Ophiolite: the Clipperton transform and the 9°17'N Deval magma chamber depth and petrological implicafrom forward modelling of common depth point tions. In: GASS, L. G., LIPPARD,S. J. & SHELTON, data. Journal of Geophysical Research, 98, 13 945A. W. (eds) Ophiolites and Oceanic lithosphere. 13 969. Geological Society, London, Special Publication KHIRATOV, N. I., LEBEDEV,E. B., DORFMAN,A. M. & 13, 71-82. BEGDASSROV, N. S. 1983. Study of process of CnmSTENSEN, N. I. 1972. Compressional and shear melting of the Kirgurich basalt by the wave wave velocities at pressures to 10 kilobars for method. Geochemica, 9, 1239-1246. basalts from the East Pacific Rise. Geophysical Journal of the Royal Astronomical Society, 28, KORMENDI, F. & DIETRICH, M. 1991. Non-linear waveform inversion of plane-wave seismograms 425-429. in stratified elastic media. Geophysics, 56, 664CHRISTESON, G. L., WILCOCK,W. S. D . & PURDY, G. 674. M. 1994. The shallow attenuation structure of the LANGMUIR, C. H., BENDER, J. F. & BATIZA, R. 1986. fast-spreading East Pacific Rise near 9°30'N. Petrological and tectonic segmentation of the East Geophysical Research Letters, 21, 321-324.

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Pacific Rise, 5°30'-14°30'N. Nature, 332, 422-429. LIN, J. & PARMENTIER, E. M. 1988. Mechanisms of lithospheric extension at mid-ocean ridges. Geo-

physical Journal of the Royal Astronomical Society, 96, 1-22. MACDONALD, K. C. & Fox, P. J. 1988. The axial summit graben and cross sectional shape of the East Pacific Rise as indicators of axial magma chambers and recent volcanic eruptions. Earth and Planetary Science Letters, 88, 119-131. SEMPERE,J.-C. & Fox, P. J. 1984. East Pacific Rise from Siqueiros to Orinoco FZ: Along strike continuity of axial neovolcanic zone and structure and evolution of overlapping spreading centres. Journal of Geophysical Research, 89, 6049-6069. , Fox, P. J., PERRAM, L. J. Er AL 1988. A new view of the mid-ocean ridge from the behaviour of ridge-axis discontinuities. Nature, 335, 217-225. MCKENz~E, D. P. 1985. The extraction of magma from the crust and mantle. Earth and Planetao' Science Letters, 74, 81-91. MANGHNANI, M. H., SATO, H. & RAI, C. S. 1986. Ultrasonic velocity and attenuation measurements on basalt melts to 1500°C: Role of composition and structure in the viscoelastic properties. Journal of Geophysical Research, 91, 9333-9342. MARSU, B. D. 1989. Magma chambers. Annual Review of Earth and Planetary Science, 17, 439-474. MAVKO, G. M. 1980. Velocity and attenuation in partially molten rocks. Journal Q[ Geophysical Research, 85, 5173-5189. MORTON, J. L., SLEEP, N. H., NORMARK, W. R. & TOMPKINS, D. H. 1987. Structure of the southern Juan de Fuca Ridge from seismic reflection records. Journal of Geophysical Research, 92, 11 345-11 353. MURASE, T. ~ MCBIRNEY, A. R. 1973. Properties of some common igneous rocks and their melts at high temperature. Geological Society of America Bulletin, 84, 3563-3592. MUTTER, J. C., CARBOTTE,S. M., Su, W. S. ETAL. 1995. Seismic images of active magma systems beneath the East Pacific Rise between 17o05' and 17°35'S Science, 268, 391-395. O'CONNELL, R. J. & BUDIANSKY,B. 1977. Viscoelastic properties of fluid-saturated cracked solids. Journal of Geophysical Research, 82, 5719-5735. PHIPPS MORGAN, J. & CHEN, Y. J. 1993. The genesis of oceanic crust: Magma injection, hydrothermal circulation, and crustal flow. Journal of Geophysical Research, 98, 6283-6297. --, HARDING,A., ORCUTT, J., KENT, G. & CHEN, Y. J. 1994. An observational and theoretical synthesis of magma chamber geometry and ,

crustal genesis along a mid-ocean ridge spreading center, In: RYAN, M. P. (ed.) Magmatic Systems, Academic Press, San Diego. RYAN, M. P. 1987. Neutral buoyancy and the mechanical evolution of magmatic systems. In: MYSEN, B. O. (ed.) Magmatic Processes." Physiochemical Principles. The Geochemical Society, Special Publication, 1, pp. 259-287. SATO, H., SACKS,I. S. & MURASE,T. 1989. The use of laboratory velocity data for estimating temperature and partial melt fraction in the low-velocity zone: Comparison with heat flow and electrical conductivity studies. Journal of Geophysical Research, 94, 5689-5704. SCHMELING, H. 1985. Numerical-models on the influence of partial melt on elastic, anelastic and electric properties of rocks, l, elasticity and anelasticity. Physics of the Earth and Planetary Interior, 41, 34-57. SINTON,J. M., ~ DETRICK,R. S. 1992. Mid-ocean ridge magma chambers. Journal of Geophysical Research, 97, 197-216. SLEEP, N. H. 1975. Formation of oceanic crust: Some thermal constraints. Journal of Geophysical Research, 80, 4037-4042. SPARKS, R. S. J. & HUPPERT, H. E. 1984. Density changes during the fractional crystallization of basaltic magmas: fluid dynamic implications. Contributions to Mineralogy and Petrology, 85, 300-309. STOLPER, E. & WALKER, D. 1980. Melt density and the average composition of basalt. Contributions to Mineralogy and Petrology, 74, 7-12. TOOMEY, D. R., PURDY, G. M., SOLOMON, S. C. & WILCOCK, W. S. D. 1990. The three-dimensional seismic velocity structure of the East Pacific Rise near latitude 9°30'N. Nature, 347, 639-645. VERA, E. E., MUTTER, J. C., BUHL, P., ETAL. 1990. The structure of 0-0.2 My old oceanic crust at 9°N on the East Pacific Rise from expanded spread profiles. Journal of Geophysical Research, 95, 15 529-15 556. WIEDICKE, M. & COLLIER, J. S. 1993. Morphotectonic characteristics of the Southern Lau basin. Journal of Geophysical Research, 98, 11 769-11 782. WORSTER, M. G., HUPPERT, H. E. & SPARKS,R. S. J. 1990. Convection and crystallisation in magma cooled from above. Earth and Planetary Science Letters, 101, 78-89. WRIGHT, D. J., HAYMON, R. M. & FORNARI, D. J. 1995. Crustal fissuring and its relationship to magmatic and hydrothermal processes on the East Pacific Rise crest (9°12 ' to 54'N). Journal of Geophysical Research, 100, 6097-6120.

Fault-controlled magma transport through the mantle lithosphere at slow-spreading ridges S. A L L E R T O N 1 & C. J. M A C L E O D 2

1Department of Geology & Geophysics, University o f Edinburgh, West Mains Road, Edinburgh EH9 3JW, UK 2 Department of Earth Sciences, University of Wales Cardiff, PO Box 914, Cardiff CF1 3 YE, UK

Abstract: At the ends of spreading segments at slow-spreading ocean ridges the axial valleys are usually asymmetric and bounded by large valley-wall faults, in contrast to segment centres, which are usually symmetric, with relatively small faults. These morphological variations are believed to reflect differences in the thermal structure of the lithosphere, caused by focused upwelling of partially molten mantle beneath segment centres. This gives rise to a thinner crust but thicker, and consequently stronger, lithosphere at segment ends, which is likely to provide a barrier to melt migration in comparison with segment centres. High-resolution sidescan sonar images of parts of the Mid-Atlantic Ridge reveal that flattopped seamounts occur preferentially at segment ends. Many of these seamounts are situated asymmetrically within the axial valley, and show a spatial association with the large segment-end faults, raising the possibility that melt may be channelled through the lithospheric mantle by such faults. If this is the case, then one might expect to encounter gabbros within mantle shear zones. We here document one such case in the Lizard Ophiolite, Cornwall, Southwest England: the Carrick Luz shear zone, a 100m wide dykelike body of gabbro and gabbro mylonite occurring within mantle ultramafic rocks. The shear zone shows a progression from penetrative mylonite fabrics, discrete ultramylonitic shear planes, through to cataclasites and fault gouges. The penetrative mylonite fabrics themselves deform, and are in turn cut by, mafic dykes. Shear direction and sense is the same for all of the fault rock types, ductile and brittle, and after correction for the regional tilt of the ophiolite section, are consistent with normal faulting. Gabbro mylonitic shear zones such as this within the lithospheric mantle provide a mechanism for weakening the lithospheric mantle. They might be expected to have a high acoustic impedance contrast with the surrounding ultramafic rocks, and are strong candidates for dipping seismic reflectors observed in old oceanic crust.

Magmatism and tectonics at slow-spreading ridges Accretion at slow-spreading mid-ocean ridges is dominated by the complex interaction of magmatism and tectonism, which vary in relative importance along-axis on the scale of the oceanic spreading segments. The slow-spreading Mid Atlantic Ridge is characterized by pronounced axial valleys, bounded by normal faults with vertical displacements of up to 2000 m. The bathymetry of the ridge axis between transform fault zones varies along strike, defining morphological segments that are typically about 50 km long and which may be offset by up to c. 20 km (Semp~r~ et al. 1990, 1993). At the centres of segments, the valley floor is usually relatively shallow, and faults are typically multiple, smallthrow structures that define a symmetric graben.

At segment ends, the valley floor is deeper, and the valley is an asymmetric half-graben, dominated by one or two large-throw faults on one side of the valley. Other slow-spreading ridges (e.g. the Southwest Indian Ridge, Patriat & Segoufin 1988) are similarly characterized by pronounced axial valleys dominated by large normal faults. Slow-spreading ridges are believed to have a relatively low magma supply, magma apparently being delivered episodically beneath the spreading centres (Whitehead et al. 1984; Sinton & Derrick 1992; Murton & Parson 1993; Sinha et al. 1997). Magma chambers are not permanent, and freeze between melt delivery episodes. Melt is largely focused at segment centres, where a thicker crustal sequence is developed. Faults play an important role in accommodating plate separation, particularly at the magma-starved segment ends (Shaw 1992; Allerton et al. 1996).

ALLERTON,S. & MACLEOD,C. J. 1998. Fault-controlled magma transport through the mantle lithosphere at slow-spreading ridges. In: MILLS,R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 29-42

29

46° 10' W

descan }

Fig. 1. (a) Geological sketch map indicating the position of discrete, flat-top seamounts on the Mid Atlantic Ridge north of the Kane Fracture Zone, observed on deepLowed sidescan sonar data. (b) Deep-towed (TOBI) sidescan sonograph from the eastern edge of the axial valley of the Mid-Atlantic Ridge at 24°20'N. Image is 3 km wide, insonified from the left (the west-southwest), and high backscatter is shown in white. A circular, flat-topped seamount approximately 100 m high is visible in the upper left portion of the image, built upon a low-reflectivity terrain of heavily sedimented, tectonized volcanics. The seamount has a central crater and is cut by several small faults on its eastern side. It lies at, and casts a shadow against, the foot of a partly talus-covered, high-backscattering fault scarp (the right-hand side of the image). This fault is a 2 km-high structure that forms the eastern valley-wall fault of the axial valley at the northern telwnination (inside corner) of the spreading segment immediately north of Lhe Kane transform fault. The much higher reflectivity and relative lack of faulting of this and other flat-topped seamounts compared to the rest of the valley floor, coupled with their spatial association with the very large valley-wall faults at the inside corners of segment ends and distinct geochemical compositions, suggests that these ~eamounts were extruded in this position and do not share the same plumbing system as the main flows of the axial volcanic ridge.

ia)

FAULT-CONTROLLED MAGMA TRANSPORT

Volcanic morphologies at slow-spreading ridges Two types of volcanic morphology predominate within the axial valley: fissure volcanoes and flattopped seamounts (Smith & Cann 1993; Smith et al. 1995; Lawson et al. 1996). Fissure volcanoes, characterized on deep-towed sidescan sonar records by linear traces of hummocky, 'cauliflower' texture volcanic material, often over linear fault scarps, are by far the most significant areally. By analogy with similar features on Iceland, and from the 1993 eruption of the Cleft segment of the Juan de Fuca Ridge, they are believed to be fed to a large extent by dyke injection laterally away from the segment centre (Embley & Chadwick 1994; Batiza 1996). Flat-topped seamounts, typically c. 1-2km across and a few hundred metres high, occur throughout the axial valley. Many seamounts, such as the 'W-Seamount' of Smith et al. (1995) are spatially associated with the axial volcanic ridge and probably represent focused effusion from a fissure (Smith & Cann 1993). Head et al. (1996), have argued that these larger volcanoes result from eruption centred at the widest parts of a feeder dyke; part of a continuum of volcanic activity which includes hummocky volcanoes and flows. Other flat-topped seamounts, however, are not associated with volcanic ridges, are often offset from the centre of the axial valley, and occur in close proximity to the large valley wall faults, especially at the inside corners of segment ends (Lawson et al. 1996). On the MAR between 24°00'N and 24°40'N, for example, we observe 14 flat-topped seamounts, of which 10 occur offset from the axis at the segment end between 24°20'N and 24°25'N (Figs la & b; Allerton et al. 1995, 1996; Lawson et al. 1996). Several lines of evidence suggest that these seamounts were extruded in this position and do not share the same plumbing system as the main flows of the axial volcanic ridge: they often appear anomalously reflective acoustically compared to the valley floor around them, indicating they are not as highly sedimented and therefore younger than the surrounding valley floor; they are less faulted than the surrounding flows (Fig. 1b) which often exhibit dense fissuring; and they are geochemically and petrologically distinct. Geochemical data from dredge hauls in the M A R N O K area, between 24°00'N and 24°40'N on the M A R (Lawson et al. 1996) show that the hummocky flows from the axial volcanic ridges of each of the two segments can all be related by simple fractional crystallisation, with MgO

31

contents decreasing progressively towards each segment end, implying that basalts in each segment evolve by fractionation within the crust from a single parent magma. In contrast, the above-mentioned flat-topped seamounts at the 24°20N discontinuity are more primitive than the surrounding fissure-fed flows and require different, more enriched parental compositions. These seamounts, therefore, have distinct melting histories and a separate plumbing system, with magmas presumably rising relatively rapidly through the lithosphere and undergoing little fractionation and homogeniZation (Lawson et al. 1996).

Strength of the oceanic lithosphere The strength of the oceanic lithosphere is controlled by its thermal structure and its lithology. Crustal rocks, modelled by diabase, yield by plastic flow at lower temperatures than ultramafic rocks (Harper 1985). The lower geothermal gradient and thinner crustal sequence postulated at segment ends (Cannat 1993, 1996; Tolstoy et al. 1993) both contribute to increasing the relative strength and thickness of the lithosphere at segment ends compared to segment centres. At segment centres melt is able to rise via cracks through the asthenospheric mantle directly into the crust; at segment ends, however, the strong, thick mantle lithosphere should act as a barrier to melts rising from the asthenosphere. The large valley-wall faults at segment ends accommodate vertical displacements of up to 2 kin. These faults are likely to cut right through the lithosphere and to extend at depth to considerable distances across the axial valley. The amount of plate motion these faults accommodate argues for protracted lifetimes. For example, Allerton et al. (1996) suggested that large faults on the M A R north of the Kane fracture zone were active for more than 0.25 Ma. Karson et al. (1987) estimated that ~0.75 Ma worth of plate separation was accommodated by extensional faulting in the M A R K area. Very large corrugated fault surfaces recently identified at the Atlantis Fracture Zone (Cann et al. 1997) may have accommodated plate separation for over 1 Ma. These observations suggest that, once developed, these faults are very much weaker than the lithosphere around them. Modelling of stresses associated with normal faults (King & Ellis 1990) predicts that the region below the fault at depth should be in relative tension, and that above the fault the hanging wall should be in compression. Shaw & Lin (1993) modelled stresses associated with a

32

S. ALLERTON & C. J. MACLEOD

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Fig. 2. Location of the Lizard Complex in Southwest England (inset), and a simplified geological map of the ophiolite, modified after Bromley (1979). Oceanic relationships are best preserved in the highest structural unit in the eastern part of the Complex, where a north- or northwest-dipping section through ultramafic tectonites, gabbros and sheeted dykes is exposed.

single large slip event on a fault typical of those encountered on the Mid-Atlantic Ridge, and indicated that much of the lower lithosphere of the hanging wall beneath the valley floor would be in relative compression. This region should therefore be expected to act as a barrier to vertical melt migration. The actual stress conditions are probably much more complicated than that represented in this simple model; individual slip events are probably smaller and more localized, and these stresses would dissipate over time. However, the conditions of stress associated with large faults can clearly be demonstrated to influence the flow of groundwaters in continental settings (Ge & Garvin 1994). Melts should be readily able to rise through tension fractures in the footwall at depth, but be inhibited in the hanging wall. The faults themselves, therefore, should provide plausible pathways for magmas rising beneath segment ends. Melts rising up such conduits may not encounter major magma chambers, and probably have a very short residence time within the lithosphere. We predict that they would thus be likely to have geochemical signatures similar to those observed in the 24°20'N flat-topped seamounts. Melt migration along faults, albeit on a smaller (metre to tens of metres) scale within the lower oceanic crustal sequence, has been documented in gabbros from the slowspreading Southwest Indian Ridge (Dick et al.

1991), in the Josephine ophiolite (Kelemen & Dick 1995) and in the Lizard ophiolite (Hopkinson & Roberts 1995). We propose that normal faults can act as pathways for melts migrating through the lithospheric mantle at slow-spreading ridge segment ends. The end product of this process should be gabbro-filled shear zones within the oceanic lithospheric mantle. Because such shear zones are unlikely to be found easily in the modern oceans with existing technology and resources, we turn to ophiolites to search for possible onland analogues.

The Lizard ophiolite We find a suitable analogue of gabbros within a shear zone in mantle rocks in the Lizard ophiolite of Cornwall (Southwest England). The Lizard Complex, which is of late Devonian age, includes a metamorphic sole, ultramafic rocks, gabbros and dolerite dyke swarms, and is associated with a m61ange containing metabasic lavas and sediment clasts (Fig. 2; Bromley 1979; Kirby 1979). The ophiolite preserves a wide range of both ductile and brittle faults, and has been interpreted as having been dismembered by shearing at a slow-spreading ridge (Gibbons & Thompson 1991; Roberts et al. 1993). In the eastern part of the ophiolite the sequence dips towards the north or northwest,

FAULT-CONTROLLED MAGMA TRANSPORT

33

Strike of ridge axis

Gabbro ~

Ultramafic rocks

I

Carrick Luz Gat Schists

Mafic dykes, showing cross-cutting relationships Faults and Shear zones parallel to ridge axis. Other faults, possibly emplacement-related. Coastline

Fig. 3. Block diagram illustrating the orientation of the principal elements of the eastern part of the Lizard Ophiolite. This figure illustrates the relative geometries of features exposed at the surface, but is not meant to reflect the deeper structure of the ophiolite, which is likely to have been sigificantly modified by emplacement tectonics. Inset: simple block diagram illustrating the present orientation of the palaeohorizontal plane (black) and a palaeovertical, ridge-parallel plane (grey). The box has the same orientation as the main figure. exposing a sequence of ultramafic tectonites, gabbros and sheeted dykes (Fig. 3). Within the gabbros north of the village of Coverack, mafic dykes, which strike north-northwest, exhibit cross-cutting relationships consistent with progressive tilting during intrusion about a gently north-northwest-plunging axis parallel to the strike of the dykes (Fig. 4a, Roberts et al. 1993). A series of minor ductile shears occur at the contact between tectonite peridotites and gabbros exposed at Coverack (Vearncombe 1980). The kinematics of these shears are consistent with normal displacement on faults which strike parallel to the dykes (Roberts et al. 1993). We also note that the crustal architecture in the Lizard ophiolite is heterogeneous in that a mafic/ultramafic cumulate sequence (e.g. Kirby 1978; Leake & Styles 1984; Gibbons & Thompson 1991) is developed only locally above the ultramafic tectonites, in the manner envisaged by Cannat (1993, 1996) for magma-starved slowspreading ridge crust. Carrick Luz shear zone

A large gabbroic mylonite shear zone crops out at Carrick Luz in coastal exposures of the southeastern part of the ophiolite, in ultramafic rocks south of Coverack (Fig. 2). This body was initially interpreted by Flett (1946) as a dyke of foliated gabbro. Gibbons & Thompson (1991)

termed it the Carrick Luz shear zone and interpreted it as a downfold of an extensional detachment fault along the petrological Moho, largely based on mapping of the coastal exposure and the published inland outcrop pattern (Flett 1946), which indicates a substantial area of gabbro to the west and north. However, ground (S. Allerton, unpublished data) and airborne magnetic surveys (Rollin 1986) and soil sampling of the area (Smith & Leake 1984) show convincingly that ultramafic rocks are much more widespread in the inland outcrop, and that the Carrick Luz gabbro is a narrow (~100m wide) north-northwest-trending, dykelike body, consistent with the coastal outcrop as recognized by Flett (1946). The body itself is gabbro and gabbro mylonite, with serpentinized peridotite on either margin, and dips moderately to steeply east-northeast. We identify a significant lithological difference between the ultramafic rocks on either side of the Carrick Luz shear zone (Fig. 5): on the southwestern side the rocks are serpentinized harzburgites and lherzolites; on the northeastern side the ultramafic rocks include troctolites and dunite pods, an assemblage characteristically encountered immediately beneath the ultramafic/mafic contact ('petrological Moho') in ODP (Ocean Drilling Program) drillholes (Gillis et al. 1993) and ophiolites (Benn et aL 1988), including the Coverack area of the Lizard (Kirby 1979).

34

S. A L L E R T O N & C. J. M A C L E O D

'In

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ua

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Coverack Ga

ea

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88

FAULT-CONTROLLED MAGMA TRANSPORT

35

Fig. 5. Cartoon section illustrating the principal features characterizing the Carrick Luz Gabbro shear zone: (1) serpentinized harzburgites and lherzolites of footwall, with crystals of pyroxene defining tectonite structure. (2) hanging wall serpentinized harzburgites and lherzolites with (3) troctolites and (4) dunite pods. (5) Gabbro dykes occur in both hanging wall and footwall. (6) Intercalated serpentinites and gabbros at the margins of the shear zone. (7) Carrick Luz gabbro with L-S tectonite fabric. (8) isolated pods of relatively undeformed gabbro, 9) discontinuous mylonite shear zones deforming L-S fabric. (10) mafic dyke deformed by L-S fabric. (11) mafic dyke cutting L-S fabric, and cut by (12) mylonite shear zones, and (13) cataclastic shear zones.

The ultramafic rocks on either side of the shear zone exhibit a high-temperature, low-strain tectonite fabric similar to that found in other ophiolites and equated to ductile flow in the asthenospheric mantle (Carter & Av6 Lallemant 1970; Nicolas & Poirier 1976). These fabrics trend at a high angle to the margin of the shear zone. Sharp-sided, but irregular gabbro dykes cut the peridotites on both sides of the shear zone margin, but are not observed within the Carrick Luz gabbro itself. At the margins of the shear zone, ultramafic rocks and gabbros are intercalated on a metre scale (Fig. 6a). Within the intercalations, the fabrics in the ultramafic rocks parallel the shear zone. Rare angular xenoliths of ultramafic material, exhibiting no sign of internal deformation, also occur within

the shear zone itself (Fig. 6b). These observations, taken together, suggest that the Carrick Luz gabbro was intruded into peridotite within the lithosphere, and that this intrusive contact has been modified by intense ductile shearing. Deformation within the shear zone is highly heterogeneous, with rare pods of almost undeformed gabbro grading into shear zones with strongly developed, often penetrative L - S - and L- type tectonite fabrics (Fig. 6c). These mylonite fabrics are cut by an undeformed mafic dyke (see fig. 4 of Gibbons & Thompson 1991) with chilled margins, which is itself offset by a narrow, dark mafic ultramylonitic shear zone. Elsewhere, an amphibolitized dolerite dyke has been sheared into parallelism with the shear zone (see Fig. 3 of Gibbons & Thompson 1991).

Fig. 4. Equal-area projections of the orientation of structures in the Eastern Lizard. (a) Mafic dykes from the gabbro section northeast of Coverack. (b) Comparison of lineation directions from L-S tectonites (crosses), discrete mylonitic shears (large circles), and cataclasite gouge zones (small dots). Orientations of shear zone lineations from Carrick Luz: in situ, and after removal of 30° and 45° tilts to the north-northwest. (e) L-S tectonites from Carrick Luz, showing orientation of foliation plane and stretching lineation direction. (d) Discrete mylonitic shear zones from Carrick Luz, showing orientation of the shear plane and the stretching lineation direction. (e) Cataclasite gouge zones from Carrick Luz, showing orientation of shear plane and slickenside or mineral fibre lineation. (f) Mylonitic shear zones from Coverack, showing orientation of shear plane and stretching lineation direction.

36

S. ALLERTON & C. J. MACLEOD (a)

(b)

Fig. 6. (a) Intercalated gabbro and peridotite at contact of Carrick Luz shear zone. Steep mylonite fabric in gabbro is visible behind and to the right of the figure; smoother-weathering surfaces in the left foreground are peridotites ('P') intercalated with mylonite gabbro ('G'). Ductile fabrics in both peridotite and gabbro are parallel to lithological contacts. (b) Angular xenolith of peridotite in mylonitic gabbro near margin of Carrick Luz shear zone.

Deformatioh shows a progression from pervasive tectonite fabrics through thinner, discrete mylonitic shear zones and felsic ultramylonites, to brittle shear zones comprised of cataclasite (sometimes foliated) and/or fault gouge (Fig. 6d). The overall direction of movement is similar for all of these fault types (Fig. 4b); plunging

gently towards the southeast, with an apparent dextral normal (top to southeast) shear sense in present-day co-ordinates (though see below). In the least-deformed portions of the Carrick Luz gabbro, igneous textures are identifiable, but plagioclase grains nevertheless display undulose extinction and closely-spaced deforma-

FAULT-CONTROLLED MAGMA TRANSPORT

37

(c)

(d)

Fig. 6. (e) Contact between mylonitic and virtually undeformed gabbro within the Carrick Luz gabbro body. (d) Fault gouge ('G') within mylonite gabbro. Gouge zone is dominated by clay minerals and haematite, yet has the same orientation and shear sense as the mylonites, implying that movement on the Carrick Luz shear zone was protracted and continued into the brittle field to significantly lower temperatures.

tion twins. Clinopyroxene is partially pseudomorphed by brown hornblende and colourless actinolite. The pervasively deformed gabbro mylonites are characterized by syn-kinematic greenish to colourless amphibole (although relict clinopyroxene does survive in some porphyroclasts), and sphene and zoisite replacing inter-

mediate-Ca plagioclase. The predominantly hydrous mineralogy of the mylonites indicates the ingress of water into the shear zone during deformation. By analogy with the assemblages described by Hopkinson & Roberts (1995) in the mylonite shear zones from Coverack, the mineralogy of the Carrick Luz mylonites is indicative

38

S. ALLERTON & C. J. MACLEOD flattopped

•

axial volcanic

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....... ....

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A"

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" """~'~,

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Fig. 7. Schematic, approximately true-scale, cross-section of a slow-spreading ridge axis near a segment termination, illustrating the relationships deduced from sidescan sonar records from the MAR and from observations of the Lizard ophiolite. Melt delivery is episodic, and insufficient to keep up with plate separation. Thick lithospheric mantle is present at the ridge axis itself. Large valley-wall faults at the inside corners of axial discontinuities extend right across the axial valley to the base of the lithosphere, and are responsible for accommodating some plate separation. At least at times of waning magma supply, the crustal sequence is thin, irregular and poorly developed in comparison to segment centres (see Cannat, 1996, for a less conservative version), and a large component of this crust is fed laterally from the segment centres. Temporal variation in melt supply causes lateral variations in crustal thickness. Melt delivered from the upwelling, cenvecting asthenosphere is delivered to the base of the lithosphere and rises by buoyancy, leaving gabbroic dykes and irregular intrusions similar to those drilled on the MAR at 23°20'N (Cannat et al. 1995). The rising melts encounter the valley-wall faults and, because of the relative compressional stresses in the hanging wall, some at least are channelled up the valley-wall faults rather than crossing them. When erupted, these melts form flat-topped seamounts located asymmetrially within the axial valley, close to the large valley-wall faults, and with melting and/or fractionation histories distinct from those lavas fed from the segment centres. The remnants of these channelized melts at depth are deformed by continuing motion on the valley-wall faults, leaving mylonitized gabbros in shear zones in lithospheric mantle peridotite. Presence of the melts probably contributes to weakening of the faults. If and when magmatic activity resumes on a larger scale, this mechanism breaks down, the lithosphere thins and the fault is cut. Relict portions of the fault zone are rafted off-axis, those on the right of the sketch dipping away from the axis. These gabbro mylonite remnants in the mantle lithosphere or lowermost crust should have extremely high impedence contrasts with the surrounding (partially serpentinised?) peridotites and are good candidates for the outward-dipping seismic reflectors observed in old MAR crust.

of d e f o r m a t i o n largely in the lower amphibolite to u p p e r greenschist facies; however, the preservation of relict h i g h - t e m p e r a t u r e phases (e.g. the rare aluminous hornblende) suggests that shearing almost certainly started at m u c h higher temperatures, but that evidence for it has been largely obliterated by c o n t i n u i n g deformation. The Carrick Luz gabbro shares some impor-

tant characteristics with large-scale continental e x t e n s i o n a l faults, specifically the a p p a r e n t vertical offset between the footwall and the h a n g i n g wall, a n d the progressive change from ductile t h r o u g h to brittle deformation. In its present orientation the Carrick Luz shear zone has a gently plunging lineation and, at face value, m i g h t be interpreted as a strike-slip

FAULT-CONTROLLED MAGMA TRANSPORT fault; however, the kinematics of the shear zone are similar to those of small-scale shears exposed at Coverack (Fig. 4f), which have been interpreted as representing ridge-parallel normal faults (Roberts et al. 1993). It is clear that the sequence has been tilted during its emplacement. Roberts et al. (1993) determined from the orientation of dykes that the eastern part of the ophiolite has been tilted by about 30 ° about an east-northeasterly trending axis. Our own measurements of dykes from the gabbroic sequence support this deduction (Fig. 4a). Restoration of this late-stage tilt returns the multiple sets of dykes to a common, northnorthwesterly strike direction. The latest dykes of mafic suite are restored to vertical by this correction, suggesting that it is an appropriate correction to return the ophiolitic sheet to its original orientation. If this tilt of 30 ° is removed, the Carrick Luz shear zone becomes an oblique/ normal structure dipping at about 50 ° to the northeast, and striking parallel to the sheeted dykes. A slightly greater tilt (~45 °) about the same axis brings the Carrick Luz shear zone to a purely normal shear sense (Figs 4c-e), with a direction of movement similar to those of the smaller shear-zones at Coverack. No other palaeoverticals or palaeohorizontals are preserved to allow the original orientation to be constrained more precisely. It is difficult to quantify precisely the total displacement associated with the Carrick Luz shear zone. The existence of near-pervasive deformation fabrics with crystal elongation ratios in excess of 10:1 across the entire 100 m width of the shear zone, suggests extremely high shear strains (Nicolas 1987). Discrete mylonites and ultramylonite bands several centimetres wide may each offset mafic dykes by several metres, implying shear strains of the order of 100 : 1. The thickest of these ultramylonite bands is about 0.5m. In addition, cataclastic gouge zones are often several tens of centimetres thick, also suggesting substantial displacements, perhaps of the order of tens of metres, or more (e.g. Scholz 1987). Altogether, we estimate that the cumulative displacement associated with these faulting styles across the Carrick Luz shear zone as a whole, is likely to be in excess of a kilometre.

Discussion The ultramafic rocks of the footwall and hanging wall to the Carrick Luz structure have asthenospheric tectonite fabrics which are cut by the shear zone, and exhibit only minor shearzone parallel deformation compared to the

39

gabbros within the shear zone. This suggests that the ultramafic rocks were relatively strong compared to the gabbro, and were within the lithosphere when intrusion and deformation of the gabbros occurred. This strength contrast could conceivably be isothermal, relying entirely on the different rheologies of mafic and ultramafic rocks (Harper 1985), or it might result from differences in temperature following the intrusion of gabbroic melt along the shear zone. The latter alternative is more consistent with the progressive evolution from highly ductile to brittle fault rocks in the shear zone, and with the observation of deformed xenoliths/lenses of ultramafic rocks within the shear zone. From the spatial position of the gabbro body, presence of undeformed or little-deformed gabbro dykes in the adjoining footwall and hanging wall, and brittle deformation of the peridotite host rocks, we conclude that the Carrick Luz gabbro body was intruded into a fault zone within the lithospheric mantle (Fig. 5). Because the shear zone is cut by an undeformed dolerite dyke, this intrusion and deformation must have occurred at the ridge axis (Fig. 7). We suggest that structures such as Carrick Luz provide important pathways for melts at the ends of slowspreading ridge segments, and also a mechanism for weakening the lithospheric mantle. Origin o f s e i s m i c reflectors

Deep seismic reflection profiles acquired in the slow-spreading oceanic lithosphere of the Atlantic Ocean (White et al. 1990; Henstock et al. 1995) have imaged reflectors in the crust and upper mantle. Some of these reflectors dip towards the axis, and correlate with a step in the basement, and can confidently be interpreted as faults formed at the spreading axis. Others, which occur only in the (seismically defined) lower crust and upper mantle, and dip away from the axis, are more enigmatic. Some of these deep reflectors represent out-of-plane scattering from basement topography (Kent et al. 1997). In other cases the reflectors have been mapped in 3D grids, and must represent in situ reflectors (J. Collier, pers. comm.) Suggestions for the origin of these reflectors includes magmatic features and hydrated shear zones. The Carrick Luz shear zone may be expected to have a high acoustic impedance contrast compared to the ultramafic rocks of the footwall and hanging wall, enhanced by anisotropy, with fast seismic velocities of minerals parallel to the shear zone. The shear zone represents an important conduit for the ingress of water into the lower crust and mantle. Its width (~100 m) is

40

S. ALLERTON & C. J. MACLEOD

approximately a quarter of the wavelength of seismic reflections recorded from deep crustal/ upper mantle depths (~400-500 m), the requisite thickness to produce constructive interference, so the shear zone should act as a strong seismic reflector (Fountain et al. 1984). At the segment end, the main normal fault accommodates the majority of the separation between the two lithospheric plates, and may therefore be considered to represent the plate boundary. If this old fault is cut when the plate boundary moves to a new locus, either a new, steep fault within the valley or an active axial volcanic zone, then the upper part of the old fault will migrate passively with one plate, and the lower part with the other. This lower part, with reflection characteristics similar to the Carrick Luz shear zone, will thus dip away from the spreading axis, and terminate in the middlelower crust. These truncated faults thus share many of the characteristics of the outwarddipping seismic reflectors. Volcanic architecture We suggest that melt rises from the asthenosphere up the large normal faults to a crustal level where formation of vertical extensional fractures is favoured over continued transport up existing normal faults. Flat-topped seamounts will be extruded over these fractures, offset from the surface trace of the large valley wall faults. This mode of accretion will tend to place the flat-topped volcanoes asymmetrically within the valley (Fig. 4), away from the axial volcanic ridge. These volcanoes are volumetrically small, although they will contribute, with talus fans and landslips, to an enhanced thickness of extrusives and erosional sediments adjacent to the largest of the fault scarps. Growth faulting within the half graben, accommodated by progressive tilting of the footwall of the large fault, would result in a highly asymmetric pattern of accretion, and the development of wedges of extrusives which thicken away from the axis (Fig. 5).

Conclusions At slow-spreading ridges, most melt delivered from mantle is focused at segment centres. The segment ends are much cooler, and a thick mantle lithosphere develops at the axis. This thick lithosphere further acts as barrier to melt transport from mantle. At segment ends, large, asymmetrical valley wall faults are important in accommodating plate separation. These faults, moderately dipping, extend across and beneath

the axial valley. Evidence from the Lizard ophiolite and from the Mid-Atlantic Ridge at 24°N suggests that the melt may be channelled by these large faults, eventually giving rise to seamounts erupted close to the faults, offset from the centre of the axial valley. Melt lubrication of such faults has the potential to significantly weaken the lithospheric mantle; in addition, the gabbro mylonites produced by deformation on the faults are likely candidates for the dipping seismic reflectors observed in old oceanic crust and (probably partially serpentinized) mantle lithosphere.

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FAULT-CONTROLLED M A G M A T R A N S P O R T evolution of an in-situ section of oceanic layer 3. Proceedings of ODP, Scientific Results, 118, College Station, TX (Ocean Drilling Program), 439-538. EMBLEY, R. W. & CHADWICK, W. W. 1994. Volcanic and hydrothermal processes associated with a recent phase of seafloor spreading at the northern Cleft segment, Juan de Fuca Ridge. Journal of Geophysical Research, 99, 4741-4760. F5ETT, J. S., 1946. Geology of the Lizard and Meneage (second edition). Geological Survey of Great Britain Memoir. FOUNTAIN, D. M., HURICH, C. A. & SMITHSON, S. B. 1984, Seismic reflectivity of mylonite zones in the crust. Geology, 12, 195-198. GE, S. M. & GARVIN, G. 1994. A theoretical model for thrust-induced deep groundwater expulsion with application to the Canadian Rocky mountains. Journal of Geophysical Research, 99, 1385213 868. GIBBONS, W. & THOMPSON, L. 1991. Ophiolitic mylonites in the Lizard Complex: ductile extension in the lower oceanic crust. Geology, 19, 10091012. GILLIS, K. M., MI~VEL, C., ALLAN, J. F. Er AL. 1993. Proceedings of ODP, Initial Reports, 147, College Station, TX (Ocean Drilling Program), 366 pp. HARPER, G. D. 1985. Tectonics of slow spreading midocean ridges and consequences of a variable depth to the brittle/ductile transition, Tectonics. 4, 395409. HEAD III, J. W., WILSON,L. & SMITH,D. K. 1996. Midocean ridge eruptive vent morphology and substructure: evidence for dike widths, eruption rates, and evolution of eruptions and axial volcanic ridges. Journal of Geophysical Research, 101, 28 265-28 280. HENSTOCK, T. J., WHITE, R. S. & MCBRIDE, J. H. 1995. The OCEAN study area: tectonic history from magnetic anomaly data and seismic reflectivity. Journal of Geophysical Research, 100, 2005920 078. HOPKINSON, L. & ROBERTS, S. 1995. Ridge axis deformation and coeval melt migration within layer 3 gabbros: evidence from the Lizard Complex, U.K. Contributions to mineralogy and petrology, 121, 126-138. KARSON, J. A., THOMPSON, G., HUMPHRIS, S. E., EDMOND, J. M., BRYAN, W. B., BROWN, J. R., WINTERS, A. T., POCKALNY,R. A., CASEY,J. F., CAMPBELL,A. C., KLINKHAMMER,G., PALMER,M. R., KINZLER, R. J. & SULANOWSKA,M. M. 1987. Along axis variations in seafloor spreading in the M A R K area. Nature, 328, 681-685. KELEMAN,P. B. & DICK, H. J. B. 1995. Focused melt flow and localized deformation in the upper mantle: Juxtaposition of replacive dunite and ductile shear zones in the Josephine peridotite, SW Oregon. Journal of Geophysical Research, 100, 423-438. KENT, G. M., DETR1CK,R. S., Swivr, S. A., CONNINS,J. A. & KIM, I. I. 1997. Evidence from Hole 504B for the origin of dipping events in oceanic crustal reflection profiles as out-of-plane scattering from

41

basement topography. Geology, 25, 131-134. KING, G. C. P. & ELLIS, M. 1990. The origin of large local uplift in extensional regions. Nature, 348, 689-693. KIRBY, G. A. 1978. Layered gabbros in the Eastern Lizard, Cornwall, and their significance. Geological Magazine, 115, 199-204. - 1979. The Lizard Complex as an ophiolite. Nature, 282, 58-61. LAWSON, K., SEARLE,R. C., PEARCE, J. A., BROWNING, P. & KEMPTON,P. 1996. Detailed volcanic geology of the M A R N O K area, Mid-Atlantic Ridge North of Kane transform. In: MACLEOD, C. J., PARSON, L. M. & WALKER, C. (eds) Tectonic,

Magmatic, Hydrothermal and Biological Segmentation of Mid-Ocean Ridges, Geological Society, London, Special Publication, 118, 61-102. LEAKE, R. C. & STYLES,M. T. 1984. Borehole sections through the Traboe hornblende schists, a cumulate complex overlying the Lizard peridotite. Journal of the Geological Society of London, 141, 41-52. MURTON, B. J. & PARSON, L. M. 1993. Segmentation, volcanism, and deformation of oblique spreading centres, a quantative study of the Reykjanes Ridge. Tectonophysics, 222, 237-257. NICOLAS, A., 1987. Principles of Rock DeJormation. Reidel, Dordrecht. - & POIRIER, A. 1976. Crystalline Plasticity and Solid State Flow in Metamorphic Rocks. Wiley, New York. PATRIAT,P. & SEGOUFIN,J. 1988. Reconstruction of the Central Indian Ocean. Tectonophysics, 155, 211234. ROBERTS, S., ANDREWS, J. R., BULL, J. M. & SANDERSON, D. J. 1993. Slow-spreading ridge-axis tectonics: evidence from the Lizard Complex, UK. Earth and Planetary Science Letters, 116, 101-112. ROLLIN, K. E. 1986. Geophysical surveys on the Lizard Complex, Cornwall. Journal of the Geological Society of London, 143, 437-446. SCHOLZ, C. H. 1987. Wear and gouge formation in brittle faulting. Geology, 15, 493-495. SEMVERE, J.-C., PURDY, G. M. & SCHOUTEN,H., 1990. Segmentation of the Mid-At/antic Ridge between 24°N and 30°40'N. Nature, 344, 427-431. , LtN, J., BROWN, H., SCHOUTEN,H. & PURDY, G.M., 1993. Segmentation and morphotectonic variations along a slow-spreading center: the MidAtlantic Ridge (24°00'N-30°40'N). Marine Geophysical Research, 15, 153-200. SHAW, P. R. 1992. Ridge segmentation, faulting and crustal thickness in the Atlantic Ocean. Nature, 358, 490-493. - & LIN, J. 1993. Causes and consequences of variations in faulting style at the Mid-Atlantic Ridge. Journal of Geophysical Research, 9 8 , 21 839-21 851. SINHA, M. C., NAVIN, D. A., MACGREGOR, L. M., CONSTABLE, S., PEIRCE, C., WHITE, A., HEINSON, G. & INGLIS, M. A. 1997. Evidence for accumulated melt beneath the slow-spreading MidAtlantic Ridge. Philosophical Transactions of the Royal Society, London, 355, 233-253.

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SINTON, J. M. & DETRICK, R. S. 1992. Mid-ocean ridge magma chambers. Journal of Geophysical Research, 97, 197-216. SMITH, D. K. & CANN, J. R., 1993. Building the crust at the Mid-Atlantic Ridge. Nature, 365, 707-715. , DOUGHERTY, M. E., LIN, J., SPENCER, S.,' MACLEOD, C. J., Keeton, J., McAllister, E., Brooks, B., Pascoe, R. & Robertson, W., 1995. Mid-Atlantic Ridge volcanism from deep-towed side-scan sonar images, 25°-29°N. Journal of Volcanology and Geothermal Research, 67, 233262. SMITH, K. & LEAKE, R. C. 1984. Geochemical soil surveys as an aid to mapping and interpretation of the Lizard Complex. Journal of the Geological Society of London, 141, 71-78.

TOLSTOY, M., HARDING, A. J. & ORCUTT, J. A. 1993. Crustal thickness on the Mid-Atlantic Ridge: bull's eye gravity anomalies and focused accretion. Science, 262, 726-729. VEARNCOMBE,J. R. 1980. The Lizard ophiolite and two phases of suboceanic deformation. In: PmqAWOTOU, A. (ed.) Proceedings of the International Ophiolite Symposium, Cyprus, 1979. Cyprus Geological Survey Department, 527-537. WHITE,R. S., DETR1CH,R. S., MUTTER,J. C., BUHL, P., MINSHULL, T. A. • MAURICE, E. 1990. New seismic images of oceanic crustal structure. Geology, 18, 462465. WHITEHEAD, J., DICK, H. & SCHOUTEN, H., 1984. A mechanism for magmatic accretion under spreading centres. Nature, 312, 146-148.

Structure, petrology and seafloor spreading tectonics of the Kizildag Ophiolite, Turkey YILDIRIM

DILEK 1 & PETER THY 2

I Department of Geology, Miami University, Oxford, OH 45056, USA 2Department of Geology, University of California, Davis, CA 95616, USA

Abstract: The Kizildag ophiolite in southern Turkey is a remnant of the Neo-Tethyan oceanic crust and displays well-preserved magmatic and tectonic structures of seafloor spreading origin. The ophiolite consists of two structurally distinct massifs that are separated by the NW-striking high-angle Tahtak6pr~ fault. The main massif to the west contains a serpentinized peridotite core adjoined on the southeast by the normal faultbounded plutonic sequence and sheeted dyke complex in a structural graben. The dykegabbro boundary within this graben is in places faulted along a low-angle detachment surface and is locally marked by a transition zone with mutual intrusive relations between the dykes and isotropic gabbros and plagiogranites. This igneous boundary contains numerous proto-dyke intrusions marking a well-preserved root zone of the sheeted dyke complex and may represent the roof of a fossil magma chamber. Mineralized oceanic faults within the dyke complex form two major subsets. Dyke-parallel normal faults form horst and graben structures and locally flatten with depth acquiring a listric geometry. Dykeperpendicular faults display steep dips and subhorizontal slickenside lineations, suggesting their oblique- to strike-slip nature. The graben structure containing the plutonic sequence and the sheeted dyke complex is analogous to those documented in the Troodos ophiolite and may similarly represent a fossil spreading axis. The second massif east of the Tahtak6prti fault consists mainly of serpentinized peridotites directly overlain by lava flows, rotated dyke blocks, and gabbros. Sulfide mineralization along some fault planes in the extrusive rocks indicates that hydrothermal systems were associated and operated synchronously with magmatic and tectonic extensional processes. Stratigraphic relations and the structural architecture in this massif suggest that the Kizildag oceanic crust underwent crustal denudation and unroofing of the upper mantle as a result of tectonic extension at a spreading centre. The Tahtak6prti fault separating the two massifs is an accommodation zone that permitted differential movements between the adjacent ridge segments during generation of the Neo-Tethyan oceanic lithosphere. The general structure of the ophiolite suggests its evolution via seafloor spreading and an asymmetric simple shear extension along a slow-spreading centre. The sheeted dykes and pillow lavas have relatively high SiO2 and AlzO 3 and low FeO and TiO2 concentrations and show limited FeO and TiO2 enrichments with decreasing MgO contents. These compositional properties differ markedly from typical tholeiitic suites from the ocean floor and arc settings, but correspond closely to those documented from the sheeted dyke complex and the lower volcanic suite of the Troodos ophiolite. The major element compositions suggest low pressure and high degree of melting from a depleted mantle source. The trace element concentrations are markedly depleted in both high-field strength and rare-earth elements and relatively enriched in largeion-lithophile elements compared to normal MORB. It is inferred that the Kizildag ophiolite formed in a tectonic setting where melting occurred at relatively low pressures within the stability field of spinel and proceeded to high melt fractions possibly by progressive depletion and melt removal. A limited positive correlation between the extent of melting and light rare-earth enrichment can be related to infiltration by fluids or melts derived from a more fertile source or possibly a subducting slab fragment. Comparison of the structure and petrology of the Troodos and Kizildag ophiolites and the regional geology suggest their evolution along a seafloor spreading system within the Southern Neo-Tethys. Spreading probably occurred in short segments and resulted in development of the Cretaceous Neo-Tethyan seaway as a marginal basin between the Tauride platform in the north and Afro-Arabia in the south.

Three fifths of the surface of the solid earth is made of oceanic lithosphere, all of which has been formed during the last 160 Ma at the mid-

ocean ridges. Understanding the structure of the oceanic lithosphere and the mid-ocean ridges is particularly important because it provides a key

Dn~EK, Y. & THY, P. 1998. Structure, petrology and seafloor spreading tectonics of the Kizildag Ophiolite, Turkey. In: MILLs, R. A. & HARRISON, K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 43-69

43

44

Y. DILEK & P. THY

Fig. 1. Location of the Kizildag ophiolite within the Neo-Tethyan ophiolite belt around the Arabian promontory in the eastern Mediterranean region (modified from Lippard et al. 1986). Key to the numbers: 1. TekirovaAntalya (Turkey), 2. Troodos (Cyprus); 3. Mersin (Turkey), 4. Kizildag (Turkey), 5. Baer-Bassit (Syria), 6. Khoy (Iran), 7. Kermanshah (Iran), 8. Neyriz (Iran), 9. Makustan (Iran), 10. Semail (Oman). to understanding the mantle and the kinematics of plate tectonics. Models dealing with the generation and development of oceanic lithosphere have been based mainly on inferences from seafloor bathymetry, interpretations of seismic reflectors based on continental analogues, mantle flow patterns, and rock samples (Macdonald & Luyendyk 1986; Orcutt 1987; Dick et al. 1991; Smith & Cann 1993). Recent geophysical studies and in situ sampling of the oceanic lithosphere by the Deep Sea Drilling Project (DSDP) and the Ocean Drilling Program (ODP) have provided new information and hypotheses on its architecture and evolution (Mutter et al. 1985; Detrick et al. 1987; Orcutt 1987; Sempere & Macdonald 1987; Toomey et al. 1990; Vera et al. 1990; Kong et al. 1992; Sinton & Detrick 1992; Alt et al. 1993; Dilek et al. 1996a). The results of these studies show that the widely accepted model of the late 1970s for the 'layered-cake' structure of oceanic crust and a steady-state magma chamber beneath the ridge axis is inadequate to explain the complex architecture of oceanic lithosphere. In slowspreading ocean ridges, magma chambers appear to be ephemeral, with long intervening periods of amagmatic (tectonic) extension producing major structural disruptions. Rotated blocks of crust and high- to low-angle normal faults and detachment surfaces are present at the

crest of the Mid-Atlantic Ridge (Karson 1990; Mutter & Karson 1992), and lower crustal rocks (undeformed to mylonitized gabbros) and serpentinized peridotites are exposed on the rift walls of various mid-ocean ridges, suggesting the exhumation of deep crust and mantle sections as a result of tectonic stretching during ocean crust generation (MARK area, Hess Deep, Site 735B at the SW Indian Ridge). Certain limitations exist, however, regarding the tectonic interpretations derived from geophysical investigations, bathymetric surveys, and indirect sampling and discontinuous coring of modern oceanic lithosphere. For example, equivocal origins of crustal reflectors due to the difficulties of tracing them to the seafloor pose problems for interpreting fault geometries and fault kinematics and for better constraining the distribution of deformation in the oceanic lithosphere. Although they have been instrumental in studying the structure and petrology of lower crust and upper mantle, dredged rock samples from seafloor outcrops are problematic because of the poorly-constrained geological setting and vertical dimensions of the rock bodies from which these samples are derived. In addition, current technology limits penetration depths of drilling and seismic resolution in young oceanic lithosphere. Ophiolites, which are good structural analogues for oceanic crust, provide 3D exposures and age relations to study the nature

45

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE ISKENDERUN......

Cover Units ~

Altuvium(Quaternary)

~

Basalt(Quaternary)

~

Sandstone,argillite,evaporites, limestone (Neogene) Umestone,mart,andsandstone

~ ~

J KIRIKNAN

(Palaeogene)

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(U. Cretaceous)

Autochtonous Ur~its ~ Limestone,marl, andsandstone Limestone,dotomiticand ootithiclimestone

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%

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Plutonicsequence(isotropicgabbro, cumulategabbro,plagiogranite) eridotites(harzburgitetectonite, dunite,serpentinite) Amphibolite olcanosedimentarysequence

~ I ~

a~/,

i..///fzz/~.ti

tltlil@J~Tii t i t i i t l ~ l l l l l l l t i l i l i i t l p i i l t i l i i i t i i V i i t J i l i i t i t i i i i l i l i t i l l ~ . l i l i i l l i l i t i . / ~ J i / / / / t / / . liiltliii. ///////i/. Itilililt. illt~illi.41t] ililtill~ "% Iif¢it11] lii illt~ illJ'til~ ,~tttliltl.

SAMAN

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Fig. 2. Geological map of the Kizildag (Turkey) and Ba~r-Bassit (Syria) ophiolites (data from Selguk 1981; Tekeli et al. 1983; Tekeli & Erendil 1983; Erendil 1984; Dilek & Delaloye 1992). of extensional deformation and magmatic construction of crust at spreading environments. Therefore, they complement our knowledge of the architecture of oceanic crust derived from both seismic images and drill holes at modern spreading centres. Intact ophiolites without polyphase deformation fabrics associated with obduction processes record significant information on the temporal interplay between structural and magmatic processes and on the role of

faulting and brittle failure in development of hydrothermal systems. The Kizildag ophiolite in southern Turkey is a remnant of the Neo-Tethyan oceanic lithosphere with a well-preserved pseudostratigraphy and displays structural evidence for seafloor spreading tectonics and associated structural, magmatic, and hydrothermal processes (Erendil 1984; Robertson 1986; Tekeli & Erendil 1986; Piskin et al. 1990; Dilek et al. 1991; Dilek &

46

Y. DILEK & P. THY

Delaloye 1992; Dilek & Eddy 1992; Lytwyn & fault; Fig. 2). The main massif to the southwest Casey 1993). It occurs in the peri-Arabian contains a NE-trending antiformal core of ophiolite belt that includes the Troodos (Cy- peridotites bounded by crustal units both on prus) and Semail (Oman) ophiolites (Fig. 1), the NW and the SE (Dilek et al. 1991). The 3which have been indispensable in the evolution km-thick peridotitic core consists mainly of of ideas concerning the origin and significance of serpentinized harzburgitic tectonite with local ophiolites and ophiolite-ocean crust analogy bands and lenses of dunite, wehrlite, lherzolite, (Moores & Vine 1971; Gass & Sinewing 1973; and feldspathic peridotites. Tectonized harzburPanayioutou 1980; Coleman 1981; Lippard et al. gites locally display a well-developed foliation 1986; Nicolas 1989; Dilek et al. 1990; Malpas et defined by mineral flattening that is commonly al. 1990). Recent structural studies in the parallel to the layering and banding in the Kizildag ophiolite have shown that its internal mantle sequence. These compositional layers/ structure is mainly a manifestation of seafloor bands are isoclinally folded, and the foliation is spreading processes, and that the spatial and generally parallel to the axial plane of the folds temporal relations between magmatic and tec- (Dilek & Eddy 1992; Dilek, unpublished data). tonic features and the areal distribution of Gabbroic to diabasic dykes crosscut the folds ophiolitic subunits are a result of extensional and foliation planes in the mantle rocks, and tectonics that the ophiolite underwent in in- they are relatively undeformed running nearly traoceanic conditions (Erendil 1984; Tekeli & parallel to the general trend of the main massif Erendil 1986; Dilek et al. 1991; Dilek & Delaloye (Dilek et al. 1991). The harzburgite and the 1992; Dilek & Eddy 1992). Systematic structural associated ultramafic lenses are composed of and petrological studies in the Kizildag ophio- olivine (Fo89 92), o r t h o p y r o x e n e (En89_ lite, therefore, complement the information and Fs08Wo03), diopside (En53Fs05Wo43), and chroknowledge gained from the Troodos and Semail mite. The plagioclase in the gabbroic lenses of ophiolites, provide important insights into mag- the tectonites is An9~ in composition. The matic and tectonic development of oceanic crust, layered gabbroic cumulates and isotropic gaband generate new constraints on the evolution of bros contain olivine (Fos~s2), clinopyroxene the Neo-Tethyan oceanic lithosphere. (EnsoFs08Wo42), and plagioclase (Ans5 89) (PiThis paper describes the internal structure of skin et al. 1990). the Kizildag ophiolite in comparison to the The contact between the mantle rocks and the structural architecture of modern oceanic litho- plutonic sequence dips away from ultramafic sphere formed at constructive plate boundaries rocks and beneath the gabbros (Fig. 3A). On the with different spreading rates, introduces a southeast flank of the mantle core, this contact is structural model for the kinematics of extension characterized by a 50-100m thick zone comassociated with seafloor spreading during its posed of highly altered and fine-grained, sheared evolution, and presents new data and interpreta- rocks, and the gabbros above display cumulate tions on its petrogenesis and tectonomagmatic textures commonly cut by low-angle mylonitic evolution. shear zones and boudinage structures (Fig. 4). Millimetric to decimetric shear zones have diffuse to sharp boundaries with the surrounding undeformed gabbros, and pyroxene grains are Internal structure stretched and elongated with asymmetric tails The NE-SW trending Kizildag ophiolite consists within the shear zones. Locally, gabbros show a of a core of serpentinized mantle rocks overlain weakly- to moderately-developed foliation, deby a plutonic sequence, sheeted dykes, and fined by the preferred orientation of acicular extrusive rocks and is stratigraphically overlain primary crystals (pyroxene and plagioclase), that by a generally east-dipping sedimentary se- is parallel to the layering. Isotropic gabbros quence consisting, from bottom to top, of become predominant upward in the plutonic conglomerate, sandstone, and neritic to pelagic sequence as the layered and foliated fabric limestones (Fig. 2; Dubertret 1953/55; (~ogulu disappears. Multiple and mutual intrusive rela1973; ~ogulu et al. 1975; Selguk 1981; Tinkler et tions between isotropic gabbros and small al. 1981; Tekeli et al. 1983; Delaloye & Wagner bodies of plagiogranite, leucocratic gabbro, 1984; Erendil 1984; Dilek & Moores 1985; Tekeli and dolerite are common in the uppermost part & Erendil 1986; Piskin et al. 1990; Dilek et al. of the plutonic sequence (Fig. 5; Dilek & Eddy 1991; Dilek & Delaloye 1992; Lytwyn & Casey 1992). The entire plutonic section is cross-cut by 1993). The ophiolite includes two structurally veins and veinlets containing amphibole, chlordistinct massifs that are separated by a NW- ite, and plagioclase. The vein density increases striking high-angle oblique fault (Tahtak6prti near the top of the isotropic gabbros where vein

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE ,

NW

Kizildag Mtn. .....

structural graben

-

47 SE

1 2 Kizildag Ophiolite Rrr~k.~wAv

No verticalexaggeration

~,,namoer

B

MARK (C) I^l

EPR East PacificRise at 9°N Latitude E ~ Extrusiverocks Breccia.teddykes I..lydrothermalflow Melt lens Plutontcrocks with layering Interlayeredmaficutlramaflcrocks Mantleuitramafics

KJzildagOphleliteand MARKArea m ~ ~ ~ ~ [~

Talusbreccia Sedimentarycover E~usive rocks Sheeteddykes Plutonicrocks Se6~ntinlzedpeddotlte with gabbrodykes

EPR at 9°N Latitude

Fig. 3. Interpretive geological cross-sections at the same scale from: (A) Kizildag ophiolite; (B & C) slowspreading oceanic lithosphere in the northern and central spreading cells MARK (N) and MARK (C), respectively, in the MARK area along the Mid-Atlantic Ridge (MAR); and (D) 0.0-0.2 million year old oceanic lithosphere along the East Pacific Rise (EPR) at 9°N latitude (no vertical exaggeration). The northern spreading cell, MARK (N), along the MAR undergoes more than 80-90% tectonic extension with a higher magma budget compared to the central spreading cell, MARK (C), further south. The area of MARK (C) experiences more than 100% tectonic extension with a low magma budget and is characterized by the existence of a serpentinite detachment surface (Cannat et al. 1995). The cross-section from the East Pacific Rise is based on expanded spread profiles and several common depth point reflection lines along and across the rise axis and on the assumption that there exists a symmetry with respect to the rise axis (Vera et al. 1990). LVZ denotes the low velocity zone that extends to a distance no greater than 10 km away from the axis. The zone of molten material at the top of LVZ depicts a narrow axial magma chamber confined to within less than 2 km from the rise axis. Note the apparent lack of deformation and faulting in the crustal units of the young oceanic lithosphere at EPR. The internal architecture of the Kizildag ophiolite is reminiscent of the intermediate stages of the evolution of the MARK (C) area. See text for further discussion.

networks in differentiated leucocratic rocks and gabbros produce locally well-developed breccia zones. Basaltic dyke intrusions increase towards the top of the plutonic sequence, and dykes become predominant with irregular and commonly diffuse boundaries with the host gabbro (Fig. 6). Further up from this zone (within 100-1 5 0 m ) gabbros diminish, and dykes become planar with sharp boundaries having one- and

two-sided chilled margins. This local zone of the dyke-gabbro transition appears to be the root zone of the sheeted dyke complex. The m a i n o u t c r o p of the sheeted dyke complex occurs in a N E - S W oriented synform bounded by inward-dipping faults overlying the plutonic sequence (Figs 2 & 3). These moderately- to gently-dipping faults are best exposed along the coastline of the Mediterranean Sea

48

Y. DILEK & P. THY

Fig. 4. Layered gabbro with mylonitic shear zones that show intense strain localization and boudinaged and necked compositional bands. This deformed layered sequence was underplated and intruded by a relatively more plagioclase-rich massive gabbro; notice the thin apophysis/dyklet emanating from this intrusion upwards into and

Fig. 5. Multiple and mutual cross-cutting relations between doleritic and plagiogranite dykes and isotropic gabbros in the uppermost part of the plutonic sequence in the ophiolite. (A) A plagiogranite dyke (Plgr) intruding into doleritic dyke swarms is in turn cross-cut by a diabasic dyke (D2) in the centre. (B) Dyke infiltration into the isotropic gabbros and magmatic stoping between the dykes and gabbros in a mush state.

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE

49

Fig. 6. Igneous dyke-gabbro contact in the ophiolite near Ikizk6prfi. (A) Near subhorizontal and gently undulating surface in the foreground is in a fine-grained isotropic gabbro, which contains irregular basaltic dykes and swirls of basalt intrusions. NE-striking, SE-dipping sheeted dykes (in the background) are emanating from this surface. (B) Close-up view of the contact surface shown in (A) Irregular basaltic dykes and dyklets with sharp steps, cusps, and segments form floating intrusions in the isotropic gabbro. The occurrence of these intrusions increases upwards near the sheeted dykes seen in A above. The lack of chilled margins along the dyke margins and the irregular shapes of the dykes suggest that the host isotropic gabbro was not completely cooled at the time of dyke emplacement into it. and are associated with anastomosing b r i t t l ~ ductile shear zones and altered rocks (Dilek & Eddy 1992). Locally, the faulted contact between the sheeted dyke complex above and the gabbroic rocks below is defined by a gentlydipping shear zone with marked strain localization, extensive brecciation, and hydrothermal veining and alteration both in the gabbroic and diabasic dyke rocks (Fig. 7). Cross-cutting relations and textural and compositional differences indicate the existence of at least three main generations of dyke intrusions in the sheeted dyke, complex. Coarse-grained, green diabasic dykes and fine- to medium-grained, gray basaltic

dykes form the first and second dyke generations, respectively, and they are commonly cross-cut by cm-scale, dark gray basaltic dykes of the third generation. First-generation diabasic dykes display N E strikes, whereas second-generation basaltic dykes commonly have N N W strikes (Figs 8A & B). The majority of the dyke rocks in all three generations are aphyric, fine- to medium-grained, subophitic-textured dolerites. Secondary replacement of the phenocryst assemblages, as well as of the groundmasses, is extensive, and the dykes commonly appear as metadolerite, composed of feldspars and green to brown actinolite with rare relict plagioclase

50

Y. DILEK & P. THY

c.±.

=

z.u

slgma

N

C.I.=2.0

sigma

N

C.I.

=

2.0

sigma

C.I.

=

2.0

sigma

Fig. 8. Lower-hemisphere equal-area contour diagrams of planar fabric elements in sheeted dykes and gabbros from the Kizildag ophiolite. (A) Poles to dyke margins in the southern part of the sheeted dyke complex (near ~evlik). Strong maxima in the SE quadrant indicates that the majority of dykes dip steeply to moderately to the northwest. (B) Poles to dyke margins in the northern part of the sheeted dyke complex (near Arifobasi and Ikizk6prfi). Strong maxima in the NW quadrant indicates that the majority of dykes dip steeply to moderately to the southeast. Thus the southeast-dipping dykes in the west and northwest-dipping dykes in the east, along with the normal faults, define a major northeast-trending structural graben in the ophiolite. The scatter in azimuth reflects a deviation in dyke orientation from first- to second-generation dykes. (C) Poles to mineralized fault planes and shear zones in the sheeted dyke complex. Two nearly perpendicular fault systems are characterized by northeast-striking and southeast or northwest dipping faults that are parallel to the mean dyke orientation and north-northwest striking and steeply northeast dipping faults. (D) Poles to mineralized fault planes and shear zones in isotropic gabbros below the sheeted dykes. Fault orientation and geometry mimic those in the sheeted dykes. Two nearly perpendicular fault systems have steep easterly dips with northeast and north-northwest azimuths. C.I. = Contour interval.

(~An77; Piskin et al. 1990) and augite grains. Mineralized oceanic faults transect the dykes and form two major subsets (Fig. 8C). One subset of faults is parallel to the dykes, has generally shallower dip angles than the dykes, and displays down-dip plunging slickenside lineations (Fig. 9A). Locally, these faults form welt-developed horst and graben structures, whereas in some places they are listric in geometry, associated with rotated and tilted fault blocks of sheeted dyke swarms (Dilek & Eddy 1992). Some of these faults are synmagmatic and appear to have facilitated and/or controlled the mode of late-stage dyke injections

into the host sheeted dyke complex (Fig. 9B). The second subset contains faults that are perpendicular to the mean dyke orientation with steep dips and oblique to subhorizontal slickenside lineations. Shear zones and faults in the isotropic gabbros below the d y k e - g a b b r o boundary mimic this fault geometry and orientation (Fig. 8D). The northeast-striking contacts between the ophiolitic subunits, sheeted dykes, and gabbroic to mantle rocks are cut and offset by a number of NW-striking tear faults (Fig. 2). These faults die out in the mantle rocks to the northwest and are overlapped and covered by Maastrichtian and younger sedimentary rocks in

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE

Fig. 7. Shear zone contact between the isotropic gabbro below and diabasic sheeted dykes above. The shear zone dips gently to the NW. Both the gabbro and the dyke rocks show intense cataclastic deformation and associated brecciation and alteration within and along the narrow shear zone.

51

Fig. 10. Pillow lavas in the Kizildag extrusive sequence are truncated and separated by a SE-dipping highangle normal fault that dies out in the underlying pillow lava horizons. Locally, it follows the chilled and bulbous pillow margins. Throw is ~1.2 m.

Fig. 9. Fault relations in the sheeted dyke complex. (A) NE-striking, SE-dipping sheeted dykes are cut and offset by more gently dipping normal faults. Notice the rotation of sheeted dykes along F2. (B) The steeply NW-dipping chilled contact between the dykes 1 and 2 (D1 and D2) is offset for ~25 cm by a NE-striking, SE-dipping fault along which the third dyke (D3) is necked and thinned. D3 is not cut by the fault, however, suggesting that faulting and dyking were contemporaneous. D1 = oldest dyke, D3 = youngest dyke in the photo.

52

Y. DILEK & P. THY

the southeast (Dilek & Delaloye 1992). Both plutonic and dyke rocks are hydrothermally altered along these faults. The second massif of the ophiolite, occurring east of the Tahtak6prti fault, lacks the coherent internal structure and stratigraphy observed in the main massif in the west. It consists mainly of volcanic, dyke, and plutonic rocks directly overlying the serpentinized peridotites (Fig. 2). The contacts between the serpentinized peridotites and crustal units are commonly faulted (Dilek & Delaloye 1992; Dilek & Eddy 1992). The extrusive sequence crops out in two main localities in the eastern massif (Fig. 2). The volcanic outcrop near the village of Tahtak6pru is ~400 m thick and includes massive and pillow-lavas intercalated with metalliferous sedimentary rocks (Erendil 1984). These volcanic rocks overlie serpentinized peridotites along a gently southeast-dipping normal fault and are in turn overlain stratigraphically by the Maastrichtian siliciclastic and carbonate rocks. In some places in this section, dykes and sills occur subparallel to pillow margins; commonly, pillow lavas display steep to vertical inclinations with trends perpendicular to moderately dipping dyke swarms that intrude them. Pillow lava flows are locally truncated by high-angle normal faults, which in places follow pillow margins and/or die out down-section in the extrusive sequence (Fig. 10). Several NW-SE striking high-angle oblique faults, parallel to the Tahtak6prfi fault, transect the extrusive sequence, and are associated with highly brecciated and bleached zones containing widespread sulfide minerals. The second locality of extrusive rocks occurs farther north around the K6mfirgukuru village (Fig. 2), where volcanic rocks overlie the isotropic gabbros along fault contacts. This section, nearly 600 m thick, includes mainly pillow and massive lava flows interstratified with metalliferous umbers. The umber horizons become both progressively darker because of Mn enrichment and more silicified in the upper levels of the extrusive sequence, and they are in places spatially associated with bleached hydrothermal alteration zones and/or mineralized zones enriched in pyrite, chalcopyrite, and malachite (Robertson 1986). The volcanic rocks in both localities are composed mainly of finegrained, dominantly augite- and plagioclasephyric pillow lavas and breccias. A group of relatively primitive pillow lavas has been referred to as 'sakalavites' by Dubertret (1953/55) and Piskin et al. (1990) and shown by Laurent et al. (1980) to contain olivine (Fo76), plagioclase (An83), and augite phenocrysts.

Comparison with modern oceanic crust The seafloor morphology, the mode and occurrence of magmatic and tectonic structures on the seafloor and in oceanic crust, and the internal architecture of oceanic lithosphere are strongly controlled by spreading rate and magma supply along mid-ocean ridge spreading centres (Carbotte & Macdonald 1994; Phipps Morgan et al. 1994). The internal architecture of fast-spreading oceanic lithosphere is significantly different from that of slow-spreading oceanic lithosphere. Seismic reflection and refraction surveys have shown a continuous magma chamber beneath much of the EPR (from 9°N to beyond 13°N), except at large axis offsets, as evidenced by a crustal low velocity zone and a largely continuous shallow reflector (Derrick et al. 1987). Multichannel seismic reflection data indicate the existence of an axial magma lens containing a large (25%) fraction of melt located about 0.6 seconds beneath the rise axis and at the base of the sheeted dyke complex along the EPR (Phipps Morgan et al. 1994). This melt lens lies 1.6 km below the seafloor at 9°N on the EPR and marks the low velocity zone (Fig. 3D). The axial magma chamber at this latitude continues for nearly 5 km at depth down to the MOHO, which is interpreted to be nearly 1 km thick consisting of interlayered mafic and ultramafic rocks (Fig. 3D; Vera et al. 1990). The crustal sequence at the EPR is lacking any significant block faulting and/or tectonic discontinuities (except transform faults). Rock units with considerably different densities and velocities do not seem to be juxtaposed across fault surfaces because such structural disruptions have not been observed in seismic reflection records. This apparent lack of normal faulting and crustal stretching results from robust magmatism and steady magma supply that keep pace with fast spreading along the EPR. In contrast to the relatively undeformed nature of oceanic crust at fast-spreading ridge environments, the modern oceanic lithosphere at slow-spreading ridges is deformed in lithospheric conditions as a result of tectonic extension and stretching during and after its magmatic construction (Crane & Ballard 1981; Purdy & Detrick 1986; Detrick et al. 1987; Karson et al. 1987; Cannat et al. 1988; Karson 1990; Mutter & Karson 1992). Teleseismic earthquake and microseismicity studies show that some normal faults rupture to depths of as much as 8 to 10 km (Toomey et al. 1988; Kong et al. 1992; Mutter & Karson 1992), across the full thickness of the crust (Figs 3B & C). The concentration of most

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE earthquakes below the walls of the axial valley of the MAR suggests that these are the sites of the deepest faulting (Mutter & Karson 1992). Drilled or sampled lower crust and mantle rocks at shallow crustal levels and/or on the seafloor away from fracture zones indicate that significant episodes of tectonic extension associated with denudation and isostatic uplifting have occurred in slow-spreading ridge environments (MAR, Tiezzi & Scott 1980; Cannat et al. 1988; K a r s o n 1990; M i d - C a y m a n Rise, CAYTROUGH 1979; Southwest Indian Ridge, Cannat et al. 1991). Some of the gabbroic rocks recovered from slow-spreading ridge segments display gneissic to mylonitic textures along lowto moderately dipping ductile shear zones characteristic of high-temperature metamorphic conditions (Mevel & Cannat 1991; Dilek et al. 1997). These observations indicate that extensional deformation in slow-spreading oceanic lithosphere occurs by both brittle and ductile mechanisms through widespread normal faulting, crustal denudation, and plastic deformation during tectonic extension in the absence of a magma chamber (Harper 1985; Dilek & Eddy 1992). Whereas magmatic extension is facilitated by the emplacement of magma by fissure eruption, dyke injection, and plutonism, tectonic extension is driven by rifting and plate separation during periods of low to no magma supply beneath the slow-spreading ridge axis. Therefore, the internal architecture of slow-spreading oceanic lithosphere is a manifestation of an interplay between magmatic and tectonic extension through periods of fluctuating magma flux to the spreading system. Several lines of evidence suggest that the observed internal structure and stratigraphy of the Kizildag ophiolite are of intraoceanic origin associated with the tectonomagmatic processes that operated during its development: (1) the tear faults that are perpendicular to the main trend of the ophiolite and to the northeast-trending sheeted dyke complex are overlapped by undeformed upper Maastrichtian sedimentary rocks, indicating that these faults must have developed prior to the deposition of these cover rocks in Late Cretaceous time; (2) tectonic imbrication and thrust faulting are virtually absent within the ophiolite although the Arabian platform carbonates are exposed in several isolated tectonic windows in the east. In addition, none of the normal fault systems documented in the ophiolite occurs in the

53

underlying tectonic basement, suggesting their formation prior to the emplacement of the ophiolite onto the leading edge of the platform. The seafloor spreading origin of the faults, rather than obduction and/or post-obduction origin, is suggested by the following: (1) they are generally mineralized (epidote + chlorite + hematite + albite 4-quartz) and show slickenside lineations with down-dip plunge (dyke-parallel normal faults); (2) they are cross-cut by basaltic dykes and hydrothermal veins; (3) faults/shear zones and associated deformed rocks between the peridotites and gabbros and the gabbros and sheeted dykes show shear sense indicators (i.e. S-C fabric, asymmetric porphyroclasts, mini-faults, necking and boudinage) compatible with their extensional origin and are locally cut by pegmatitic gabbro, basaltic dykes, and/ or hydrothermal veins; (4) high-angle faults locally form horst-graben structures within the sheeted dyke complex; (5) faults in the sheeted dyke complex and those separating the dykes and gabbros, and gabbros and peridotites collectively form a structural graben, which is an extensional feature; and (6) dyke-parallel normal faults in the sheeted dyke complex die out in the uppermost part of the underlying gabbros. Points 1 through 3 above indicate that the observed faults and shear zones are of extensional origin formed in intraoceanic conditions while magmatism and associated hydrothermal circulation were still active. Most of the dykeparallel, high-angle normal faults contain lowto medium-temperature hydrothermal minerals, indicating that these brittle structures played a significant role for fluid flow and hydrothermal circulation when magmatism was still active. Points 4 and 5 also suggest a spreading-related origin of faults, inconsistent with a contractional style associated with convergence and ophiolite emplacement. Intraoceanic extensional normal faults are also ubiquitous within the sheeted dyke complexes of the Troodos and Josephine (California) ophiolites. Dietrich & Spencer (1993) have demonstrated that tectonic extension within the upper crustal levels of the Troodos ophiolite was accommodated by a combination of large dykeparallel faults, sub-horizontal shear zones, and a dense network of small-scale normal faults. The presence of cross-cutting dykes and epidote mineralization in the fault gouges indicates that

54

Y. DILEK & P. THY

N

Fig. 11. Fault and slickenside orientation from the sheeted dyke complex along the (~evlik-Arsuz coastline. Majority of the normal faults have northeast strikes with down-dip plunging slickenside lineations (poles on the great circles). Thick arrows depict the approximate orientation of the least principal stress (03) as 325° .

the observed extensional faults and shear zones in the sheeted dyke complex of the Troodos ophiolite formed within the plate accretion zone. In the Josephine ophiolite, locally multiply rotated sheeted dykes, 50° tilting of dykes and gabbro relative to the overlying sedimentary rocks, growth faults in the pillow lava flows, and massive sulphide deposits associated with the normal faults are collectively interpreted to have resulted from high degrees of crustal extension at the palaeo-spreading axis (Alexander & Harper 1992, and references therein) and are analogous to those observed in the Kizildag ophiolite. Structures documented in the Kizildag ophiolite resemble features observed in oceanic crust at modern slow-spreading centres and spreading centre-transform fault intersections (Dilek et al. 1991; Dilek & Eddy 1992). The well-developed sheeted dyke complex in the ophiolite suggests extensive magmatic extension. The northeasttrending block-faulted synform in the main massif that contains the plutonic and dyke rocks is reminiscent of median valleys of slow-spreading ridge segments (Macdonald 1982; Dilek & Delaloye 1992). The inward-dipping, dyke-parallel normal faults define a graben structure that is similar to the structural grabens (Solea and

Larnaca grabens) documented in the Troodos ophiolite (Varga & Moores 1985; Dilek et al. 1990) and to the spreading axes of slowspreading mid-ocean ridges (Macdonald 1986). Extensive normal faulting within the sheeted dyke complex indicates that magmatic extension was accompanied and/or followed by brittle deformation as a result of tectonic extension. The slickenside data combined with the fault geometry show the general orientation of the least compressive stress (0-3) to be ~325 ° (Fig. 11), which is nearly perpendicular to the mean trend of the inferred fossil spreading axis and of the main massif. The faulted contacts between the mantle sequence and plutonic rocks and between the gabbros and sheeted dyke complex suggest that these rheological boundaries were decoupled during periods of tectonic extension. The presence of hydrothermal veins, dykes and sills along and across these fault contacts and associated shear zones indicate that faulting was followed by renewed magmatism in intraoceanic conditions, rather than occurring during the emplacement of the ophiolite. Shear-sense indicators (i.e. S-C fabric, asymmetric porphyroclasts, mini-faults, necking and boudinage) in faulted and deformed gabbros near and within the fault zones give normal-sense of shearing compatible with the kinematics of extensional deformation. Mylonitic discrete shear zones and locally well-developed foliation planes in cumulate to isotropic gabbros are reminiscent of those observed in drilled core samples from gabbros recovered in the M A R K area (Mid-Atlantic Ridge near the Kane Transform; M6vel & Cannat 1991; Cannat et al. 1995; Dilek et al. 1997) and from Site 735B at the Southwest Indian Ridge (Cannat et al. 1991), and indicate that these lower crustal rocks underwent plastic deformation under intermediate- to high-temperature deformation in lithospheric conditions. The uniformly thin plutonic sequence in the second massif east of the Tahtak6prfi fault suggests a substantial amount of attenuation and crustal thinning in this part of the ophiolite. Isolated blocks of sheeted dykes occur as normal fault-bounded slivers overlying the sheared gabbros northwest of the K6mfir~ukuru Village. Different orientation of sheeted dykes in these fault slivers (Fig. 2) suggests block rotation associated with extension during crustal thinning and stretching. Brecciated, faulted, and tilted pillow lavas directly overlie the serpentinized peridotites in the same massif (Fig. 2). These observations suggest that upper crustal levels of the ophiolite were partly to entirely stripped away due to tectonic extension and that

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE NW

55 SE

Spreading

A

B

C

Kizildag Mtn. ~

'

a

~

S

~

e

C

s

structural

graben

,z/,] O

I

5kin

WESTERN MASSIF

I

No vertical exaggeration ,IL ,

"

'

,Ik

~"

................. ' -

-

Allochthonous basaltic lavas ~.~..,y',,.,,,,~.~On~ukum ........

E

Basattlc dykes ~,__ ,,,_ ,_. ^ DCl=,dlUU idvcu~ ...... n'inih-

,gl~, EASTERN MASSIF

Fig. 12. SeafIoor spreading and asymmetric simple-shear tectonic extension model for the Kizildag ophiolite shown on sequential cross-sections. (A) Seafloor spreading and associated magmatic extension along eastnortheast-trending ridge segment (in present coordinates), developing the Kizildag ophiolitic crust. (B) Asymmetric tectonic extension along a low-angle master normal fault following magmatic construction of the ophiolite. Block-faulting in the hanging wall of this detachment surface caused thinning of crustal units. Vertical narrow dome beneath the spreading axis represents a postulated late-stage ephemeral magma chamber. (C) Domal uplift of serpentinized peridotites as a result of continued tectonic extension and isostatic rebound of unloaded footwall, and extensive serpentinization and associated diapirism. Warping and partial erosion of the detachment surface occurred at advanced stages of tectonic denudation. (D) In the massif west of the Tahtak6prfi fault (accommodation zone), stripping away of crustal units and further uplift of the mantle rocks were accompanied and/or followed by renewed magmatism, which resulted in the emplacement of small gabbroic plutons and diabasic dykes into the serpentinized peridotites. (E) In the massif east of the Tahtak6prti fault, renewed magmatism following the tectonic extension episode resulted in the emplacement of small gabbroic intrusions and diabasic dykes into the uplifted serpentinized peridotites and extrusion of basaltic pillow lavas on the serpentinites already exposed on the seafloor. See text for further discussion.

lower crustal levels and upper mantle rocks were exposed on the seafloor during and/or prior to the eruption of pillow lavas (Dilek et al. 1991). Thus, the late-stage magmatic activity in the ophiolite appears to have followed a period of crustal denudation and unroofing of the mantle

rocks as a result of amagmatic extension. Oblique-slip faults that are normal to the contacts and to the inferred palaeo-spreading axis in the main massif of the ophiolite are analogous both in size and geometry to the transfer faults on the eastern wall of the M A R in

56

Y. DILEK & P. THY

the Trans-Atlantic geotraverse program (TAG) area near latitude 26°N (Karson & Rona 1990). Transfer faults in modern oceanic crust of slowspreading origin are generally oblique-slip in nature and facilitate differential extension and rotation between adjacent blocks along the spreading axis (Karson & Rona 1990). The oblique-slip faults in Kizildag are pre-Maastrichtian in age, as inferred from the existence of the Maastrichtian and younger strata unconformably overlying them, and are subparallel to the inferred spreading direction (~325 ° in present coordinate system) as deduced from the mean dyke orientation. The two massifs separated by the Tahtak6prfi fault have distinctly different internal structures. The relatively complete pseudostratigraphy and the NE-trending structural graben in the main massif to the west is juxtaposed to the east along this fault zone against a much thinner second massif, which is composed mainly of upper crustal rocks resting on the serpentinized peridotites as isolated blocks. The Tahtak6prti fault thus separates two significantly dissimilar levels of oceanic crust that display different thickness and internal architecture and is interpreted as an accommodation zone (ridge segmentation) linking two independently extending rift segments with variable extensional strain at different scales. Such accommodation zones are integral components of both oceanic and continental extensional systems and are regions of complex strike-slip and/or oblique-slip faulting (Karson 1991).

Structural evolution and tectonic model Several characteristic features help us distinguish the Kizildag ophiolite from other well-preserved ophiolites in the region (Troodos, Semail): (1) The main massif west of the Tahtak6prti fault shows an acute asymmetry with respect to the distribution of ophiolitic subunits, and the topography of its stratigraphic order is inverted. Nearly two third of the massif is made of serpentinized peridotite in an antiformal core that currently occupies the highest elevation within the ophiolite, whereas the stratigraphically higher crustal units (i.e. gabbros and sheeted dykes) occur in a graben structure in a topographic low on the southeast flank of the antiformal core. (2) The boundary between the serpentinized mantle sequence and the crustal units is mostly tectonic, dipping away from the

peridotite core. (3) Crustal units in the second massif east of the Tahtak6prfi fault are highly attenuated and thinned, juxtaposing them against each other along low-angle tectonic contacts. Extrusive rocks in this massif directly overlie the serpentinized peridotites, which are in turn cross-cut by diabasic dykes. These features of the ophiolite point to large magnitudes of tectonic extension, which might have been driven by an asymmetric extension along a low-angle detachment surface (e.g. Wernicke 1985). In this model, the Cretaceous oceanic lithosphere preserved in the Kizildag ophiolite is interpreted to have developed along a NE-trending (in present coordinate system) spreading centre and to have undergone simple shear deformation associated with asymmetric tectonic extension along a southeast-dipping low-angle fault (Fig. 12). This master fault is currently presented by the relict detachment surfaces existing on both sides of the peridotite core (Figs 2 & 3). A major part of this inferred master fault has been eroded away as a result of domal uplifting of the serpentinized peridotite (Fig. 12). The current domal shape of the postulated low-angle normal fault and the topographic high occupied by the serpentinized peridotite probably resulted from uplifting of the unloaded footwall, analogous to metamorphic core complexes in the western US Cordillera, and isostatic rebound and warping of the low-angle normal fault; uplifting of upper mantle rocks as a result of extensive serpentinization and diapiric activity during and after displacement of the oceanic lithosphere from its original spreading environment further facilitated the development of a mantle antiform (Dilek et al. 1991). This scenario is reminiscent of the tectonics of the western median valley wall along the Mid-Atlantic Ridge in the area south of the Kane fracture zone (Fig. 3; Cannat et al. 1988; Karson 1990), where mafic and ultramafic rocks occur in the footwalls of low-angle detachment faults dipping east, beneath the median valley axis. The tectonic model presented here thus infers a slow-spreading origin of the Neo-Tethyan oceanic lithosphere as exposed in the Kizildag ophiolite. These observations and interpretations suggest that oceanic lithosphere produced at a spreading centre may become tectonized and deformed within the plate boundary zone, much like continental lithosphere in continental rift zones, and that deformation fabric in some ophiolites may hence be an artefact of ocean floor tectonics rather than obduction processes.

S T R U C T U R E & PETROLOGY OF THE K I Z I L D A G OPHIOLITE

57

Table 1. Major and trace element compositions of representative dykes, gabbros, and lavas Major elements (XRF, wt%) Sample

SiO2

TiO2

A1203

94kd-02 5 4 . 0 1 0.397 16.13 94kd-ll 52.82 0.743 15.96 94kd-12 53.53 0.954 15.57 94kd-17 57.98 1.064 15.91 94kd-18a 71.35 0.708 14.45 94kd-20 52.37 0.422 15.98 94kd-21 52.13 0.838 15.38 94kd-23 52.36 0.370 13.19 94kd-25 51.04 0.310 15.93 94kd-27 68.84 0.737 13.56 Trace Elements (XRF, ppm) Sample 94kd-02 94kd-ll 94kd-12 94kd-17 94kd-18a 94kd-20 94kd-21 94kd-23 94kd-25 94kd-27

FeO

MnO

MgO

CaO

Na20

K20

P205

Total

7.44 9.39 8.95 7.68 3.73 5.27 9.61 8.06 6.05 5.35

0.146 0.137 0.119 0.061 0.019 0.113 0.164 0.152 0.120 0.032

7.17 8.71 7.28 6.09 2.01 10.09 6.85 11,23 10.76 2.51

6.15 6.82 6.75 5.03 3.23 12.22 8.53 8.49 14.23 3.24

6.96 3.73 4.60 5.00 4.62 3.16 5.01 4.12 1.57 5.21

0.19 0.18 0.32 0.30 0.26 0.18 0.01 0.11 0.08 0.40

0.026 0.055 0.075 0.082 0.189 0.040 0.060 0.030 0.027 0.198

98.62 98.55 98.15 99.20 100.57 99.85 98.58 98.11 100.12 100.08

Ni

Cr

Sc

V

Ba

Rb

Sr

Zr

Y

Nb

Ga

Cu

Zn

58 34 32 30 12 121 78 213 168 19

81 42 16 48 0 399 278 633 598 7

32 35 35 32 19 45 28 38 42 17

215 270 279 351 34 211 315 230 165 155

0 0 12 1 4 0 0 0 0 0

3 2 2 1 3 1 0 1 2 3

67 97 116 117 128 88 75 79 74 113

24 36 48 59 120 27 38 20 22 95

11 19 23 23 38 14 24 11 10 23

4 4 3 8 18 1 2 1 2 3

9 19 12 18 16 11 22 12 14 14

130 11 3 2 3 ll 7 100 88 2

53 32 14 6 5 26 46 57 34 5

Trace Elements (ICP-MS, ppm) Sample 94kd-02 94kd-11 94kd-12 94kd-17 94kd-18 94kd-20 94kd-21 94kd-23 94kd-25 94kd-27

Lu

Ba

Th

Nb

Y

Hf

Ta

U

Pb

Rb

Cs

Sr

Sc

0.19 0.30 0.36 0.39 0.63 0.23 0.39 0.22 0.16 0.36

17 14 24 8 27 19 1 10 9 9

0.20 0.15 0.18 0.98 2.72 0.18 0.23 0.13 0.09 0.31

1.79 0.90 1.08 5.14 14.69 1.09 1.39 1.30 0.55 2.39

11 18 23 23 39 14 23 12 10 22

0.58 0.95 1.34 1.63 3.18 0.75 1.12 0.52 0.53 3.08

0.12 0.07 0.08 0.35 0.96 0.07 0.10 0.10 0.04 0.17

0.04 0.06 0.08 0.23 0.68 0.07 0.08 0.05 0.03 0.21

0.76 0.15 0.21 0.19 0.22 0.42 0.46 0.64 0.29 0.19

2.7 1.2 2.0 2.2 2.2 1.4 0.7 0.9 1.l 1.9

0.52 0.01 0.01 0.01 0.01 0.01 -

92 90 111 110 125 85 79 80 76 108

39 38 36 32 15 40 39 42 44 16

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

1.48 3.17 4.40 4.54 10.95 2.36 3.55 1.44 1.58 3.54

0.70 1.50 2.02 1.88 3.96 1.13 1.68 0.71 0.74 1.59

0.33 0.62 0.77 0.74 1.69 0.51 0.76 0.30 0.34 0.69

1.22 2.31 2.99 2.75 5.30 1.77 2.63 1.26 1.22 2.63

0.26 0.46 0.58 0.57 1.02 0.38 0.58 0.27 0.26 0.50

1.82 3.18 3.99 3.80 6.76 2.37 3.94 1.94 1.75 3.46

0.41 0.71 0.87 0.88 1.47 0.55 0.89 0.46 0.39 0.79

1.22 2.06 2.49 2.50 4.18 1.56 2.53 1.39 1.13 2.27

0.18 0.30 0.37 0.38 0.63 0.23 0.38 0.21 0.16 0.34

1.16 1.86 2.26 2.41 3.97 1.46 2.37 1.34 1.07 2.17

Rare-Earth Elements (ICP-MS, ppm) Sample

La

Ce

94kd-02 1.09 2.05 94kd-ll 1.37 3.19 94kd-12 1.70 4.39 94kd-17 4,20 7.49 94kd-18a 10.31 19.69 94kd-20 1.34 2.68 94kd-21 1.74 3.73 94kd-23 1.03 1 . 7 3 94kd-25 1.04 1 . 6 7 94kd-27 1.67 3.89

Pr 0.30 0.56 0.77 0.94 2.37 0.49 0.66 0.31 0.37 0.61

Kizildag Ophiolite Samples (36°E & 36°30'N): 94kd-02 Basaltic pillow lava; 94kd-11 Sheeted diabase dyke; 94kd-12 Sheeted diabase dyke; 94kd-17 Diabasic dyke intruding gabbro; 94kd-18a Quartz diorite in the plutonic sequence; 94kd-20 Gabbroic dyke in the plutonic section; 94kd-21 Pyroxene phyric dyke intruding layered gabbro; 94kd-23 Basaltic dyke intruding gabbro; 94kd25 Isotropic gabbro in the plutonic section; 94kd-27 Quartz andesite dyke in the sheeted dyke complex.

58

Y. DILEK & P. THY

Petrogenesis of Kizildag magmas Representative suites of pillow lavas, sheeted dykes, and microgabbros were analysed by Xray fluorescence (XRF) and inductively-coupled plasma mass-spectrometry (ICP-MS) in order to evaluate the petrogenesis and interrelationships of the various ophiolitic units. Results for major and trace elements are presented in Table 1. The analysed basaltic rocks are hypersthene and olivine normative, subalkalic basalts and basaltic andesites with Mg/(Mg + Fe 2+) ratios between 0.62 and 0.80 (Table 1). Additional compositional data can be found in Laurent et aL (1980), Delaloye et al. (1979), Piskin et al. (1990), and Lytwyn & Casey (1993). The relatively unaltered dyke and volcanic rocks display restricted variations with SiO2 content between 51 and 54 wt% and MgO between 12 and 7 wt% (Table l; Fig. 13), but a small group of dykes has andesitic, dacitic, and rhyolitic compositions with SiO2 above 58 wt%. The extensive albitization of the feldspars is reflected in the low CaO and high Na20 contents shown by the majority of the analysed rocks, except for the most magnesian augite-phyric dykes (Fig. 13). Despite the effects of secondary replacements, the basaltic dykes have higher SiO2 and A1203 and lower FeO and TiO2 contents compared to typical oceanic basalt suites (Icelandic basalts) and are also less evolved with MgO content up to 12% than typical oceanic island arc suites (Fiji basalts; Fig. 13). However, the analysed Kizildag dykes compare very closely to the compositional variations found in the volcanic suites of the Troodos ophiolite (Robinson et al. 1983; Thy & Xenophontos 1991; Thy & Esbensen 1993). The relatively high A1203 contents of the basaltic rocks do not suggest extensive crystal fractionation at likely crustal conditions since such pressure ranges (10kbar) would have resulted in plagioclase saturation, and depletion in A1203 and enrichment in FeO. The relatively high A1203 and low FeO contents of the Kizildag rocks are, therefore, likely to be primary features inherited from liquids in equilibrium with a mantle peridotite. High pressure melting experiments on peridotites predict that high SiO2 and low FeO liquids may result from either low pressure or low extent of melting (Jaques & Green 1980; Takahashi et al. 1993; Niu & Batiza 1991; Langmuir et al. 1992). Low extent of melting is not so well supported by the relatively low concentrations of TiO2, as many trace elements that suggest high melt fractions or melting of a depleted source. To evaluate these possibilities,

we use the trace element concentrations, in particular those of the rare-earth elements, to estimate the nature of melting and the peridotite source that produced the Kizildag volcanic and dyke rocks. The light rare-earth (LRE) elements in the basaltic rocks are relatively depleted compared to the heavy rare-earth (HRE) elements (Fig. 14A). In addition, the basaltic rocks, compared to normal MORB, are markedly depleted in both high-field strength (HFS) and REE elements and relatively enriched in large-ionlithophile (LIL) elements (Ba, Rb, K, and Sr) (Fig. 14B). As observed for the major elements, many of the absolute and relative trace element concentrations of the Kizildag rocks show strong similarities with those of the extrusive sequences from the Troodos ophiolite (Figs 14A & C; Cameron 1985; Rautenschlein et al. 1985; Taylor & Nesbitt 1988). Both ophiolites have low incompatible trace element concentrations with strong depletion in LRE elements. The analysed rocks from the Kizildag ophiolite all display low La/SmN ratios (0.5-1.0, N--normalized to chrondite) and Dy/YbN ratios just above unity (1.1-1.2). Two groups of basaltic dykes can be distinguished on the basis of their trace element compositions: (1) a relatively low La/SmN group (0.5-0.7) with low MgO (7-9 wt%) and high TiO2 (0.7-1.0 wt%), and (2) a high La/SmN group (0.8-1.0) with high MgO (9-11 wt%) and low TiO2 (0.3-0.4 wt%). We use the REE ratios and the absolute and relative HFS element concentrations to evaluate the nature and extent of source melting inferred from the major element concentrations and to understand the two compositional groups observed among the analysed dyke rocks. Fig. 15 illustrates aggregated fractional melting models assuming a low to intermediate pressure, spinel lherzolite source. The first model utilizes a fertile primitive mantle source composition (Sun & McDonough 1989) and is shown in steps of 1% melting until 15% of melting is reached (see figure caption for details of the calculations). It is evident that this melting model fails to reproduce the observed very low Ti values even after 15% melting (just prior to melting-out of diopside). In addition, the resultant La/SmN ratio is too high to account for the origin of the low-La/SmN group (high MgO), although the estimated Ti/Zr ratios are appropriate. Another set of modelling assumes prior depletion of the source by a basaltic component (1 and 3% batch melting; DM1 and DM3, respectively) and produces a better correlation with the observed values for the low-La/SmN group but not for the

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high-La/SmN group. However, both models predict a high extent of melting (exceeding 15% for all basaltic rocks) near harzburgite melting (diopside-out). We have tested harzburgite melting by depleting a primitive mantle composition to 15% by batch melting, and subsequently melted this source assuming fractional melting. The result obtained from 1% to 15% melting illustrates the very strong depletion of Ti and La/SmN ratios not shown by the Kizildag dykes. The effects of fractional crystallization on the rare earth element ratios are less than the size of the symbols; only for the absolute concentrations of Ti will fractionation have an effect. The model illustrating fractional crystallization in Fig. 15 is shown in steps of 5% fractionation until 30% fractionation is reached and suggests that the high-La/SmN group cannot produce the low-La/SmN group, despite the fact that the major element concentrations

are generally consistent with such a model (~20% crystallization). In general, both dyke groups record a high extent of melting with the high MgO group reflecting the highest extent. The group recording the highest extent of melting also shows the least depletion in the L R E E content (highest La/ SmN ratios). We have tested the possibility that this 'positive' correlation between the extent of melting and L R E E enrichment can be related to source metasomatic transformation and modifications. The model shown and labelled E D M gives the predicted evolution of liquids extracted from a depleted mantle source (3% batch depletion of a primitive mantle) metasomatized by 1% fluid and released from a subducting oceanic slab. The fluid composition that we use reflects the composition inferred from the Mariana Trough magmas and is taken from Stolper & Newman (1994). Although our model

62

Y. DILEK & P. THY

calculation does not match the high-La/SmN group, it shows clearly that a combination of a relatively high extent of melting (20%) and a small enrichment (1%) of a slightly depleted source may in general explain the observed variation. The nature and origin of the enriching fluid or liquid is uncertain (Tatsumi et al. 1986; Taylor & Nesbitt 1988; Pearce & Parkinson 1993; Cousens et al. 1994; Stolper & Newman 1994), permitting a high degree of freedom in the model development. We infer that the Kizildag ophiolite formed in a tectonic setting where enriched fluid or liquid was derived from a deeper and more fertile source (a subducting slab?) and was added to the overlying mantle. Melting occurred at relatively low pressures within the stability field of spinel and proceeded to high melt fractions possibly by progressive depletion and melt removal. The most refractory liquids (highest MgO and lowest TiO2) are related to an enriched source, whereas the least refractory liquids (lowest MgO and highest TiO2) record little or no enrichments. Lytwyn & Casey (1993) suggested that the Kizildag ophiolite formed within a forearc environment in which fertile mantle from deeper sources migrated into and mixed with the upwelling mantle diapirs in a melting column above the Benioff zone. They envisioned that highly enriched liquids from a lower fertile source were mixed with variably refractory melts that were produced within this melting column.

Discussion on the evolution of the Kizildag ophiolite The evolution of the Troodos ophiolite (and other ophiolites in the region) in a forearc, arc, and backarc setting in a suprasubduction zone environment has been suggested previously based solely on geochemical studies of its intrusive and extrusive rocks (i.e. Miyashiro 1973; Pearce 1975; McCulloch & Cameron 1983; Robinson et al. 1983; Rautenschlein et al. 1985; Thy et al. 1985; Lytwyn & Casey 1993). Several tectonic models combining the geochemical and structural data and observations have also been put forward to explain the controversial nature of the evolution of the Troodos ophiolite. Some of these models include an Andaman Sea-type seafloor spreading system with short rift segments above a highly oblique subduction zone (Moores et al. 1984); a marginal basin evolved over an active or recently active, south-dipping Palaeo-Tethyan subduction zone (Dilek et al. 1990); a forearc setting (Mariana-Bonin-type)

above a north-dipping, young intraoceanic subduction zone (Robertson 1990). The 'juvenile arc' model of Lytwyn & Casey (1993) for the Troodos and Kizildag ophiolites is similar to the Mariana-Bonin-type forearc model, and both models assume that a pre-existing oceanic crust of an older age (Late Triassic; Robertson 1990) was rifted apart to form an extensional embryonic arc in the Late Cretaceous (~Turonian). The juvenile nature of the arc and the short time span of the inferred subduction zone are invoked in these models to explain the absence of calcalkaline volcanic and pyroclastic rocks overlying the ophiolites that are typically found in modern mature arcs. The quandary with this interpretation is, however, the lack of supporting evidence from the region for the occurrence of pre-Cretaceous oceanic crust, which is envisioned to have been subducted in the lower plate and rifted apart as the country rock in the upper plate to form the Troodos and Kizildag embryonic arcs. Stratigraphic and palaeontological evidence from volcanic and sedimentary rocks of passive margin sequences in the eastern Mediterranean region indicate episodic rift magmatism in pulses in Mid- to Late Triassic, Late Jurassic and Early Cretaceous times as part of the continental rifting along the northwestern periphery of Gondwanaland during the breakup of Pangaea (Dilek 1994; and references therein); but, the generation of oceanic crust as a continuum of continental rifting apparently did not occur until much later. The regional stratigraphic record shows that both the rift basins and the surrounding continental margins underwent substantial subsidence and associated deep marine sedimentation in the Early to Late Cretaceous, corresponding to the onset of seafloor spreading and ocean crust formation in the Southern Neo-Tethys (Dilek 1994). Thus it appears that the Troodos, Kizildag and other Tauride ophiolites in southern Turkey represent the remnants of the Neo-Tethyan oceanic crust, which developed as the continental rifting finally gave way into seafloor spreading during the advanced stages of the Gondwanaland's fragmentation, and therefore the Cretaceous Troodos and Kizildag oceanic crust had already existed prior to the initiation of any intraoceanic subduction zone in the basin. The Kizildag ophiolite shows many similarities to the Troodos ophiolite both in terms of the major and trace element variations of its rocks. Both ophiolites show signs of progressive melting of a variable depleted source (Cameron 1985; Lytwyn & Casey 1993). Slightly to markedly U-shaped LREE patterns are most strongly developed by highly refractory melts of

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE the Troodos magmas (Type II and III of Cameron (1985); Fig. 14C), but also recorded by less refractory melts of both the Troodos and the Kizildag ophiolites (Fig. 14A & C). This characteristic LREE pattern can be related to metasomatic transformations of the source by an incompatible element enriched, slab-derived component (McCulloch & Cameron 1983; Cameron 1985; Rautenschlein et al. 1985; Taylor & Nesbitt 1988). Taylor & Nesbitt (1988) modelled the Troodos REE patterns and suggested an addition of only 0.2-0.3% enriched liquid to a highly refractory source. The introduction of incompatible element-enriched fluid (or melt) might have facilitated enhanced degrees of melting or melting of relatively refractory mantle material (Taylor & Nesbitt 1988). However, the LREE enrichment and the general depletion pattern are not as strong in the Kizildag compared to the upper volcanic sequence (Upper Pillow Lavas) in the Troodos ophiolite, suggesting that the palaeo-oceanic crust preserved in Kizildag might have escaped the later 'thermal burst' that caused prolonged and higher extent of melting at Troodos. One of the critical questions in addressing the magmatic evolution of both the Kizildag and Troodos ophiolites is the nature and source of the 'wet component', which was added to the peridotite and in turn resulted in lowering the melting temperature and causing additional melting (for the depleted source). How and when was this component introduced? Is an active subduction zone dipping beneath the displaced oceanic crust (future ophiolites) a necessity for the derivation of incompatibleelement enriched fluid to hydrate the overlying mantle and hence to facilitate enhanced degrees of melting? Duncan & Green (1987) stated that highly refractory melts with MgO-rich and TiO2-poor compositions can occur in a variety of tectonic environments including mid-ocean spreading ridges, backarc basins and the forearc region of island arcs. Diapiric rising of these melts within a mantle melting column and successive melt extractions from these rising diapirs leave behind a more depleted source at shallower depths and lower pressures, generating melts compositionally similar to the Upper Pillow Lavas in the Troodos. Alternatively, the emplacement of an upwelling, hot asthenosphere (slab window) in the lower plate as a result of ridge collapse and the subsequent subduction of a ridge flank may also facilitate melting of refractory peridotite and derivation of LREEenriched liquids added to the melting column (Robertson & Dixon 1984). However, this model implies the subduction of a pre-Cretaceous

63

oceanic crust and a diachronous crustal buildup within the Troodos and Kizildag ophiolites, occurrences which are not supported by the geological data and observations. Thus the source of the fluid in the evolution of the ophiolitic magmas is still enigmatic, and the recognition of the necessity of having a hydrous component and a refractory mantle source is insufficient by itself to constrain the tectonomagmatic evolution of the Kizildag and Troodos ophiolites. Conversely, the structural architecture of both the Kizildag and Troodos ophiolites strongly suggests their generation at an oceanic spreading centre. The internal structure of the Kizildag ophiolite is reminiscent of modern oceanic crust developed at slow-spreading mid-ocean ridges and might have resulted from an asymmetric mode of tectonic extension by simple shear deformation in a slow-spreading centre, analogous to the tectonic models suggested for the M A R K area and the SW Indian Ridge (Karson 1990; Cannat et al. 1991). Similarly, the Troodos ophiolite is interpreted to have formed at a spreading centre-transform fault intersection (Gass 1968; Moores & Vine 1971; Murton & Gass 1986), where the mode and nature of interplay between magmatic accretion and tectonic extension strongly controlled the evolution of its crustal architecture (Allerton & Vine 1991; Dilek et al. 1990; Varga & Moores 1985). However, these interpretations do not necessarily call for a mid-ocean ridge setting of the ophiolites. It is important to note that spreading centres in modern backarc and marginal basins (i.e. Lau and Woodlark) have recognizable or inferred symmetric magnetic stripes analogous to the mid-ocean ridge axes, and that those with moderate to fast-spreading rates display seismically defined axial melt bodies (Morton & Sleep 1985; Collier & Sinha 1990; Baker et al. 1996). Therefore, the morphological and spreading rate-dependent characteristics of axial melt bodies and the resultant thermal structure should apply to suprasubduction as well as mid-oceanic spreading centres, although geochemical characteristics may differ (Purdy et al. 1992; Lin & Phipps Morgan 1992). On the basis of the regional geological constraints and the occurrence of the seafloor spreading structures, we envisage that the Cretaceous oceanic crust as preserved in the Kizildag and Troodos ophiolites evolved at a spreading system within the Southern NeoTethys (Fig. 16). Spreading might have occurred in short segments and propagated eastwards between the Tauride platform and Afro-Arabia, much like the spreading system in the Woodlark

64

Y. DILEK & P. THY

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Fig. 16. Simplified palaeogeographic reconstruction of the Cretaceous Neo-Tethys and its environs between Eurasia in the north and Afro-Arabia in the south (data from Dercourt et al. 1986; Dilek & Moores 1990; Dilek et al. 1990; Fourcade et al. 1993). Dotted line represents modern coast lines for reference (i.e. Adriatic, Black, Caspian, and Mediterranean Seas). BB = Ba6r-Bassit ophiolite, CACC = Central Anatolian crystalline complex, KD = Kizildag ophiolite, SC -- Sakarya continent, TO = Troodos ophiolite. The northern seaways of the wedgeshaped Neo-Tethys were undergoing contraction and closing (like a zip fastener) from west to east, leaving behind the Alpine, Apennine-Ligurian, Albanian, and Hellenide ophiolites as remnants of the Jurassic-Early Cretaceous Neo-Tethyan oceanic crust. The Northern Neo-Tethys south of the Cimmerian continent was closing at the north-dipping subduction zone (site of the future Izmir-Ankara-Erzincan suture zone). Northward convergence and subsequent collision of CACC with this subduction zone facilitated emplacement of the ophiolites derived from the Northern Neo-Tethys onto its margins. The Inner-Tauride and Southern Neo-Tethys seaways had still active seafloor spreading ~95 Ma, forming the eastern Mediterranean ophiolites. The Troodos (TO), Kizildag (KD), and Bafir-Bassit (BB) ophiolites were developing at a spreading system with short rift segments within the Southern Neo-Tethys. The eastward-propagating spreading system was separating the Arabian promontory in the south from the Tauride carbonate platform in the north.

STRUCTURE & PETROLOGY OF THE KIZILDAG OPHIOLITE basin extending westwards into the P a p u a n peninsula and rifting the W o o d l a r k and Pocklington rises (Taylor et al. 1995). This model thus suggests the evolution of the Southern NeoTethys as a small ocean basin, marginal to A f r o A r a b i a , t h a t o p e n e d up b e t w e e n the two continental fragments during the Mesozoic. Timing and kinematics of emplacement of this seafloor spreading system in a suprasubduction zone setting within this marginal basin and the source and nature of the hydrous c o m p o n e n t required for the arc chemistry of the ophiolitic rocks remain to be c o n s t r a i n e d with m o r e systematic and detailed geochemical and geochronological studies in the Kizildag and other eastern Mediterranean ophiolites.

This study is part of our ongoing project in the Kizildag ophiolite and has benefited from discussions with O. Tekeli, M. Erendil, (). Piskin, G. Harper, J. Karson, and J. Casey on the geology and petrology of this ophiolite. Y. D. would like to thank the shipboard scientists of ODP Legs 148 and 153 for lively discussions on the mid-ocean ridge tectonics and the structure of modern oceanic lithosphere which helped in distillation of the ideas presented here. Financial support for this study is provided by an NSF grant (EAR-9219064) and a Faculty Research (CFR) grant from Miami University to Y. D. and is gratefully acknowledged. Constructive reviews by three anonymous referees improved the paper.

References ALLERTON, S. ~¢ VINE, F. J. 1991. Spreading evolution of the Troodos ophiolite, Cyprus. Geology, 19, 637-640. ALEXANDER, R. J. 8¢ HARPER, G. D. 1992. The Josephine ophiolite: an ancient analogue for slowto intermediate-spreading oceanic ridge. In: PARSON, L. M., MURTON, B. J. & BROWNING,P. (eds), Ophiolites and their Modern Oceanic Analogues. Geological Society, London, Special Publications, 60, 3-38. ALT, J. C., KINOSH1TA,H., STOKKING,L. B., et al. 1993. Proceedings of the Ocean Drilling Program, Initial Reports, 148, College Station, TX (Ocean Drilling Program). BAKER, M. B. & STOLPER, E. M. 1994. Determing the composition of high-pressure mantle melts using diamond aggregates. Geochimica et Cosmochimica Acta, 58, 2811-2827. BAKER, N., FRYER, P. & MART1NEX, F. 1996. Rifting history of the northern Mariana Trough: SeaMARC II and seismic reflection surveys. Journal of Geophysical Research, 101, 11 427-11 455. CAMERON, W. E. 1985. Petrology and origin of primitive lavas from the Troodos ophiolite,

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Is the oceanic Moho a serpentinization front? T. A. M I N S H U L L ,

M. R. M U L L E R ,

C. J. R O B I N S O N , R. S. W H I T E & M . J. B I C K L E

Bullard Laboratories, Department o f Earth Sciences, University o f Cambridge, Madingley Road, Cambridge CB3 0EZ, U K Abstract" Our main constraint on the volume of melt coming out of the mantle at mid-ocean ridges is the thickness of oceanic crust determined by seismic methods. The recovery of serpentinized peridotites in the rift valleys of slow-spreading ridges has led recently to a revival of the suggestion that the seismically defined lower oceanic crust may include partially serpentinized peridotite. Both seismic and geochemical data from a recent experiment on the very slow-spreading Southwest Indian Ridge indicate that, away from the fracture zone valley, a total melt thickness of ~4 km is being generated. However, in one region, around Ocean Drilling Program (ODP) hole 735B, where ~2 km of the upper crust is estimated to have been removed, the seismic data indicate a crustal thickness of 5 km. Here, the seismically determined Moho is interpreted as a serpentinization front. Serpentinization is facilitated if there are readily available pathways for seawater to reach the upper mantle, and if the upper mantle is cool. These conditions frequently are met at fracture zones on slow-spreading ridges, in the ocean-continent transition zone at non-volcanic rifted margins, and at the axes of extinct rifts. In all these locations, serpentinization sufficient to lower mantle velocities to normal oceanic crustal velocities does not reach a depth of more than 5 km beneath the seabed, and commonly normal mantle velocities of ~8 km s-a are also reached at shallow depths. Therefore, although in some anomalous locations the seismically defined crustal thickness can only be used as an upper limit on the melt supply, for normal oceanic crust, where the seismically defined crustal thickness is ~7 km, the Moho is probably not a serpentinization front, but rather a petrological boundary between mafic rocks above and ultramafic rocks below.

The ophiolite based-model of oceanic crustal structure consists of an upper 1-2km layer of basaltic lavas and dykes, underlain by a 3-5 km layer of gabbros, beneath which is upper mantle peridotite (e.g. Penrose Conference on Ophiolites 1972). Commonly oceanic seismic layer 2 is interpreted as lavas and dykes, layer 3 as gabbros, and the Moho is thought to mark the transition from gabbro to peridotite (Cann 1974; Kempner & Gettrust 1982). This model has been highly successful at fast- and intermediatespreading ridges, with only minor modifications to the interpretation of the upper part of layer 3 based on O D P drilling at Hole 504B, where this part of layer 3 consists of sheeted dykes (Detrick et al. 1994). The main argument against the alternative model of Hess (1962), that layer 3 consisted of partially serpentinized peridotite, has been the remarkable uniformity of layer 3 velocities noted by Raitt (1963), a property which has stood the test of time (White et al. 1992). However, the recovery of serpentinized peridotites, often intruded by gabbros, by dredging, submersible diving and ODP drilling in the rift valley of the Mid-Atlantic Ridge and other slowspreading ridges has led recently to an alter-

native model for crust created at these ridges in which the lower crust consists of partially serpentinized peridotite intruded by gabbroic plutons (Cannat 1993, 1996; Fig. 1). In Cannat's model, the Moho still represents the base of the zone in which gabbroic plutons are present, but if serpentinized peridotites are considered to constitute a major component of oceanic layer 3, then the base of this layer could equally well mark the boundary between serpentinized and unaltered mantle, i.e. a serpentinization front. Our main constraint on the volume of melt coming out of the mantle at mid-ocean ridges, and hence on models of melting at ridges, is the thickness of the oceanic crust determined by seismic methods and, to a lesser extent, by gravity methods. Additional constraints come from the composition of basalts erupted at ridges (White et al. 1992), and from the composition of abyssal peridotites interpreted to be the residue of mid-ocean ridge melting (Niu & H6kinian 1997). White et al. (1992) found a good agreement between seismic crustal thicknesses and melt thicknesses inferred from rare earth element compositions of mid-ocean ridge basalts (Fig. 2). Langmuir et al. (1992) found a good correlation between seismic crustal

MINSHULL,T. A., MULLER,M. R., ROBINSON,C. J., WHITE, R. S. & BICKLE,M. J. 1998. Is the oceanic Moho a serpentinization front? In. MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 71-80

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Fig. 1. Four possible geological models for oceanic layer 3, corresponding to various interpretations for the oceanic Moho (modified from Cannat 1993). (a) Layer 3 consists entirely of gabbro, and the Moho is a transition from gabbro to peridotite (Penrose Conference on Ophiolites 1972). (b) Layer 3 consists of partially serpentinized peridotite, and the Moho is a serpentinization front (Hess 1962). (e) Layer 3 consists of gabbroic intrusions within mantle peridotites, which may be partially serpentinized (Cannat 1993); the Moho marks the base of the gabbroic layer, but not all of the material above the Moho has formed by the cooling of melt. (d) As (c), but the lower part of layer 3 consists of partially serpentinized peridotite and the Moho is a serpentinization front. 40-

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Fig. 2. Histograms of oceanic crustal thickness (away from fracture zones and hotspot influences) determined by synthetic seismogram modelling and by rare earth element inversion (uses data from White et al. 1992).

thickness a n d Nas values ( s o d i u m c o n t e n t corrected for fractionation to 8% magnesium) (Fig. 3). Such a correlation is expected if the seismic crustal thickness reflects the degree of mantle melting (Klein & Langmuir 1987; Langmuir e t al. 1992; Bown & White 1994). However, there remain large uncertainties in m a n y of the parameters involved in geochemical modelling, and the above correlations are limited by the fact that in very few locations are there coincident

SERPENTINIZATION IN OCEANIC CRUST wide-angle seismic experiments and full geochemical analyses. Hence, while these correlations lend support to the idea that the seismic crustal thickness is a measure of mid-ocean ridge melt production, some uncertainty remains. If the Moho at slow-spreading ridges is commonly a serpentinization front, the melting models may need to be re-examined.

Serpentinization at the Atlantis II Fracture Zone One problem with Langmuir et al.'s (1992) compilation (Fig. 3) is that there are few places in the world's oceans where both a modern seismic experiment and a comprehensive analysis of the geochemistry of erupted basalts have been conducted. This was one objective of a seismic and dredging experiment we conducted in 1994 aboard RRS Discovery, at the Atlantis II Fracture Zone on the very slow-spreading Southwest Indian Ridge (Fig. 4; Muller et al. 1997). Here the full spreading rate has been 1 6 m m a -1 since 11 Ma (Dick et al. 1991a), a rate sufficiently slow that most thermal models would predict a reduction in the amount of melt generated by mantle decompression, and hence an anomalously thin magmatic crust (Reid & Jackson 1981; Bown & White 1994). Our seismic experiment was centred on O D P Hole 735B, which drilled 500m of gabbro on an uplifted block on the transverse ridge of the fracture zone (Dick et al. 1991b). F o r our geochemical analyses we required basalt, which was unlikely to be present on the uplifted block around the drill site. Therefore we dredged on similar aged crust on the opposite side of the spreading centre to site 735B (Fig. 4, Site 3), where the upper crust is likely to have been formed at a similar time and from approximately the same melting column as the crust at site 735B, and in the rift valley of the Southwest Indian Ridge immediately east of the Atlantis II Fracture Zone (Fig. 4, Sites 4-9). A wide-angle seismic profile along the transverse ridge of the fracture zone showed that the seismically defined crustal thickness varied between 4 k m away from the uplifted block around the drill site and c. 5 km beneath the centre of this block (Fig. 5; Muller et al. 1997). Dredged basalts both from the rift valley and from the site conjugate to 735B showed major element compositions consistent with unusually low degrees of melting in this area (Robinson et al. 1996), with Nas values consistent with a melt thickness of 3 + 1 km according to Langmuir et al.'s (1992) correlation (Muller et al. 1997).

73

Inversion of the rare earth elements using the method of McKenzie & O'Nions (1991) gives a melt thickness of 3 + 1.5 km for the conjugate site (Muller et al. 1997). These results may be interpreted to indicate that, away from the uplifted block around the drill site, most or all of the seismically defined crust is magmatic in origin (i.e. model (a) or (c) from Fig. 1). However, around the drill site, where locally ~2 km of crust have been removed by detachment faulting (Vanko & Stakes 1991), and the seismic model (Fig. 5) indicates that on average 1 km of the upper crustal section is missing, the simplest interpretation of the seismic velocity model is that the lower 2-3 km of the seismically defined crust consists of partially serpentinized peridotite (i.e. model (d) from Fig. 1, or something similar; Muller et al. 1997). The P-wave velocity in this part of the crust is 6.9 k m s -1 (Fig. 5), corresponding to a serpentine content of 35 4-10% (Miller & Christensen 1997). The alternative interpretation, that the ridge was generating 4 k m of melt before and since 1.1.5 Ma (the approximate crustal age at the drill site), but switched to generating 6-7 km of melt for 1-2 Ma around this time, that our dredging failed to recover any samples from this melting event, and that this section of thicker crust happened to be the one that was tectonically unroofed, seems rather less plausible. In the region around the drill site, the seismic velocities in the lower part of the seismically defined crust are typical of oceanic layer 3, and the Moho generates wide-angle reflections (Muller et al. 1997). Without our geochemical data and knowledge of the variations in crustal structure along the transverse ridge, this type of structure would normally be interpreted as a gabbroic lower crust. The presence of Moho reflections indicates a sharp increase in velocity, so there is an abrupt transition from ~35% serpentinization to very little or no serpentinization ~5 km beneath the Site 735B.

Geophysical properties of serpentinized peridotite From the above discussion, we can conclude that we have one good example for a seismic layer 3 and Moho which should not be interpreted in the conventional way. Unfortunately, gabbro and partially serpentinized peridotite are not always readily distinguished using geophysical data. Certainly, both can have similar Pwave velocities. Using velocity data from oceanic crustal samples and from ophiolites, Horen et al. (1996) argue that no clear lithological

74

T. A. M I N S H U L L E T A L .

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SERPENTINIZATION IN OCEANIC CRUST S

75

735B 14

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Fig. 5. Seismic velocity model along the transverse ridge of the Atlantis II fracture zone on the Southwest Indian Ridge (after Muller et al. 1997). The velocity contour interval is 0.3 km s-1 . Numbers in water column mark OBH locations, vertical line marks ODP hole 735B, and velocity contours are annotated in km s 1. Thick lines mark the regions where the Moho depth is constrained by wide-angle reflections. The region of layer 3 marked 'LOWER' is interpreted to consist of partially serpentinized mantle, so that for a horizontal distance of c. 40 km beneath the drill site, the Moho is interpreted to be a serpentinization front.

interpretation for oceanic layer 3 can be made even if S-wave velocity information is available. However, Carlson & Miller (1997) suggest that the seismic velocity measurements are less scattered if ophiolite samples are excluded, and that the observed seismic properties of oceanic layer 3 are not consistent with those of partially serpentinized peridotite. Nevertheless, these authors still show a significant overlap in properties where P-wave velocities are in the region of 7.0 km s-1. Some distinction might be made on the basis of anisotropy. Gabbro is expected to have low degrees of anisotropy, though olivine-rich gabbros may have significant anisotropy due to alignment of olivine crystals (Christensen 1978). Serpentinite, however, may exhibit as much as 30% anisotropy for both P- and S-waves, with the slow direction normal to foliation planes (Kern et al. 1997). If there were a preferred orientation of foliation planes in the serpentinized upper mantle, and this orientation was near to vertical, such anisotropy might be detectable by seismic experiments. Unfortunately, seismic experimental geometries are rarely appropriate for detailed constraints on the anisotropic properties of the lower crust and upper mantle, though clearly experiments could be designed to focus on this issue. Alternative approaches might be to use gravity data, which places constraints on crustal densities, or magnetic data, which constrains magnetizations. Unfortunately, gravity data are

also of little help in resolving gabbro from partially serpentinized peridotite, since there is overlap in the velocity-density systematics at values typical of oceanic layer 3 (Miller & Christensen 1997). Magnetic data may have a significant contribution to make, since serpentinization produces magnetite in the lower crust and upper mantle, and this effect may be responsible for the anomalous skewness of marine magnetic anomalies at slow spreading rates (Dyment et al. 1997). However, it is not clear that magnetic measurements have sufficient resolution to resolve the effects of serpentinization from other factors which determine the magnetisation of the oceanic crust, such as the composition of erupted basalt and the degree to which it is hydrothermally altered.

Conditions for serpentinization Serpentinization of the upper mantle is facilitated if there are readily available pathways for seawater to reach the upper mantle, if the magmatic crust is thin, and if the upper mantle is cool. Laboratory measurements on antigorite indicate that it is stable to ~500°C at 200 MPa (Evans et al. 1976). However, at lower temperatures lizardite and chrysotile are the stable polymorphs and these phases predominate in oceanic serpentinites. Serpentinization temperatures, based on oxygen isotope fractionation between serpentine and magnetite, appear to be as low as ,-~100°C (O'Hanley 1996, p. 160).

76

T. A. MINSHULL E T A L .

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Fig. 6. Velocity structure of oceanic crust in various tectonic settings. Dashed vertical lines indicate velocities of 7.2 kms ~, the approximate upper limit for normal oceanic crustal velocity, and 8.0km s-1, the approximate velocity in the slow direction for unaltered mantle (Bibee & Shor 1976). (a) Oceanic fracture zones. Where authors give one-dimensional velocity-depth sections, these are used; where only two-dimensional models are given, velocity~tepth sections are taken at the bottom of the fracture zone valley (for profiles crossing a fracture zone) or at the centre of the profile (for profiles along the fracture zone valley). Data are from the Kane (Detrick & Purdy 1980; Cormier et al. 1984), Vema (Ludwig & Rabinowitz 1980; Detrick et al. 1982; Potts et al. 1986a), Oceanographer (Ambos & Hussong 1986; Sinha & Louden 1986), Charlie-Gibbs (Whitmarsh & Calvert 1986), Tydeman (Potts et al. 1986b), Blake Spur (Minshull et al. 1991), and 'Northern' and 'Central' (Henstock et al. 1996) fracture zones. (b) Velocity structures for the ocean-continent transition zones of non-volcanic passive margins. Only regions where anomalously high velocities have been detected at the base of the crust are included. Data are from the west Iberian margin (Whitmarsh et al. 1990, 1996; Pinheiro et al. 1992), the Newfoundland margin (Reid 1994), and the conjugate margins of the Labrador Sea (Chian & Louden 1994; Chian et al. 1995). Where one-dimensionalmodels are published, these are used; where only two-dimensional models are published, one-dimensional sections are taken beneath instrument locations. (e) Extinct oceanic rifts. Data are from the Labrador Sea (OBS G of Osier & Louden 1992) and from the Aegir rift in the Norwegian Sea (line V, OBH 39 and 41 of Grevemeyer et al. 1997).

Below ,,~100°C, both the reaction rate and the rate of diffusion of water through the rock mass are too slow for pervasive serpentinization to occur (MacDonald & Fyfe 1985). The conditions for serpentinization are most readily met at very slow-spreading ridges, where tectonic processes appear to dominate, seismogenic faulting penetrates the entire crust (Huang & Solomon 1988; Wolfe e t al. 1995), the upper mantle is cooled conductively, and locally hydrothermal cooling might reach the base of the crust. There is ample evidence, from direct sampling of serpentinized ultramafics and from the chemical composition of h y d r o t h e r m a l fluids, that serpentinization takes place at slow spreading ridges both close to and away from segment boundaries (Rona e t al. 1987, 1992). Palmer (1996) has speculated that hydrothermal circulation at the TAG site on the Mid-Atlantic Ridge penetrates the upper mantle. Sites such as the transverse ridge of the Atlantis II Fracture Zone, where large-scale faulting is evident, are

amongst the most favourable for pervasive serpentinization to occur at deep crustal and upper mantle levels. Yet even here, in 11.5 Ma of crustal evolution, pervasive serpentinization to the degree required to give a normal oceanic layer 3 velocity (~35%) has reached a maximum of only 5 km depth beneath the drill site, and may have only reached these levels due to the fracturing associated with detachment of the upper crust. Elsewhere on the transverse ridge, where there is no evidence for such a detachment, it has reached 4 km depth or less.

Serpentinization in other tectonic settings The above observation led us to look at other areas where seismic experiments have been interpreted to indicate serpentinization at the base of the seismically defined crust. One such tectonic setting is in fracture zone valleys at slow-spreading ridges. The M o h o in m a n y North Atlantic fracture zone valleys is inter-

SERPENTINIZATION IN OCEANIC CRUST

77

preted to be a serpentinization front (Detrick et tilted fault blocks capped by pre-rift sediments al. 1993). The interpretation is clearest when remains controversial (Whitmarsh & Sawyer velocities in the range 7.2-7.6kms -1 are ob1996). However, in several locations around served at the base of the crust (Minshull et al. the North Atlantic, the crustal section is inferred 1991), since such velocities are above the normal to be thin or non-existent (whether it is oceanic upper limit for oceanic Layer 3, but too low for or stretched and intruded continental material), unaltered upper mantle. However, much lower and serpentinized peridotites have been dredged velocities, indicating higher degrees of serpenti- and drilled on the west Iberian margin (Whitnization, have also been interpreted as due to marsh & Sawyer 1996). Most seismic velocity serpentinization, where for example the seismic models of rifted margins are published as twovelocity structure differs greatly from that of dimensional cross-sections, so selection of velonormal oceanic crust (Detrick et al. 1982). A city-depth profiles necessarily involved some compilation of fracture zone seismic velocity degree of subjectivity. The profiles compiled in structures (Fig. 6a) shows that velocities and Fig. 6c are from locations at the centre the crustal thicknesses vary markedly. N o r m a l transition zones on the west Iberian, Newfoundupper mantle velocities are reached at depths land and east and west Labrador Sea margins, varying from 1.5 to 7kin, or in some experi- where crustal thinning is greatest and seismic ments are not detected at all. Depressed mantle velocity profiles are commonly interpreted to velocities may indicate that seawater penetrates indicate serpentinization at the base of the crust. to depths as great as 8 km in the lithosphere in Here again, velocities above 7 . 2 k m s -1 are sufficient quantities to cause at least a few always reached within the upper 5 km of the percent serpentinization. However, in all cases lithosphere. the velocities reach values in excess of 7.2 km s-1 The above tectonic settings are likely to be within the upper 5.7 km of the lithosphere, and those where pervasive, high degree serpentinizawithin the upper 5km for a large majority. tion reaches the largest depth in the lithosphere, Hence in fracture zone tectonic settings also, due to the low mantle temperatures and the serpentinization sufficient to lower mantle velo- pathways for water circulation made available cities to normal oceanic crustal velocities (less • by faulting. In none of these locations does than ~ 7 . 2 k m s 1) does not reach a depth of serpentinization reach the degree required to more than ~5 km. reduce mantle velocities below 7.2 km s-1 beyond An extreme example of a slow-spreading ridge 6 km depth, and rarely beyond 5 km depth. This is the case where spreading ceases completely. is probably due to a dramatic reduction in Here again, as magmatism ceases the upper permeability at this depth: laboratory measuremantle is cooled, and the final stages of ments of acoustic velocities in peridotites suggest extension are taken up on faults. There are that crack closure occurs at about 150 MPa rather few well-constrained seismic velocity (Nicolas 1989, p. 260), corresponding to a depth structures from extinct rifts. The extinct rift in of 4 - 5 k m in the crust beneath a spreading the Labrador Sea (Osler & Louden 1992) and centre. Hence, although in many of these the extinct Aegir Rift (Grevemeyer et al. 1997) locations, the Moho is best interpreted as a both show similar velocity structures: a thin serpentinization front, such an interpretation crustal section is underlain by anomalously low does not appear to be viable for normal oceanic mantle velocities, interpreted to be a result of crust formed away from fracture zones and nonserpentinization. However, low degrees of ser- volcanic margins and at full spreading rates in pentinization (less than 10%) are inferred at excess of ~20 mm a -1, where the seismic crustal both rifts (Louden et al. 1996; Grevemeyer et al. thickness is typically ~7 km and the velocity at 1997), and even in this extreme tectonic setting the base of the crust is typically ~7.2 km s 1. Nor velocities above 7.2 km s-1 are reached within the is an interpretation of the Moho as a serpentitop 6 km of the lithosphere (Fig. 6b). nization front likely to be valid for thickened A third tectonic setting where the Moho has oceanic crust formed in the vicinity of hotspots. been interpreted as a serpentinization front is at A similar conclusion was reached by Nicolas the continent-ocean transition of non-volcanic (1989) on the basis of ophiolite studies. Nicolas rifted margins. Here again, the mantle is colder identifies two types of ophiolite: the harzburgite than normal because of vertical conductive heat type, where the crustal section is thick (~6 km) loss, if the rifting is slow, and because of and is underlain by depleted mantle which is horizontal heat loss to the adjacent continents serpentinized only around fault zones, and the (Bown & White 1995). The nature of the crust in lherzolite type, where a much thinner crustal the transition zone between crust which has clear section is underlain by pervasively serpentinized seafloor-spreading magnetic anomalies and fertile mantle. It is often difficult to distinguish

78

T.A. MINSHULL ET AL.

whether the serpentinization occurred at or near the ridge crest, or as the lithosphere was d i s m e m b e r e d during ophiolite emplacement. However, based on cross-cutting field relationships between dykes and serpentinites, C o u l t o n et al. (1995) conclude that serpentinization of the Josephine ophiolite occurred at the spreading centre, a n d that the p a l a e o - M o h o in this ophiolite is a serpentinization front. This ophiolite does not appear to fit into the categories proposed by Nicolas (1989), because the peridotire section consists of harzburgite, and except in the pervasively serpentinized ultramafic cumulates at the base of the crust, serpentinization is restricted to fault zones which are surrounded by fresh peridotite, yet the reconstructed stratigraphy indicates a magmatic section that is only 2 3 k m thick. However, these observations indicate again that high degrees of serpentinization do not reach m o r e than a few kilometres into the lithosphere.

serpentinization front, and the geophysically d e t e r m i n e d crustal thickness is a reliable i n d i c a t o r o f m e l t p r o d u c t i o n , though an approximate one since some p e r i d o t i t e m a y be p r e s e n t w i t h i n the geophysically defined crustal section.

We thank all who sailed with us on RRS D&covery cruise 208 for their support at sea, H. Dick for encouraging us to explore a variety of interpretations for the crustal structure at Hole 735B, and two anonymous reviewers for their comments. This work was partially supported by the Natural Environment Research Council through grant GR3/8838 and by a studentship to C JR. TAM was supported by a Royal Society University Research Fellowship and MRM by a Carl and Emily Fuchs Foundation Overseas Scholarship and a CVCP Overseas Research Studentship. Department of Earth Sciences, Cambridge contribution 5043. References

Conclusions

F r o m the results of our seismic and geochemical work at the Atlantis II fracture zone, together with a compilation of seismic velocity structures for locations where the magmatic crust is thin and there are likely to be readily available pathways for sea water to interact with upper mantle peridotites, we conclude the following: (1) Gabbros and partially serpentinized peridotites are difficult to distinguish geophysicaUy when P-wave velocities are around 7 . 0 k m s 1, as is c o m m o n l y found in the lower part of oceanic layer 3. (2) T h e M o h o m a y be i n t e r p r e t e d as a serpentinization front where high seismic velocities (greater than ,-~7.2 k m s 1) are present at the base of the crust, or where independent constraints are available on the crustal thickness. Such an interpretation is plausible in m a n y fracture zones, at non-volcanic passive margins, in extinct rifts, and in crust formed at very slowspreading ridges such as the Southwest Indian Ridge. However, in all these locations, serpentinization sufficient to reduce velocities to normal crustal velocities only reached the upper ~ 5 k m of the lithosphere. Here, the geophysically determined crustal thickness (i.e. the crust above the M o h o ) is probably an upper b o u n d on the melt thickness. (3) In normal oceanic crust, where the M o h o is at ~ 7 k m depth, or in thickened oceanic crust, the M o h o is unlikely to be a

AMBOS,E. L. & HUSSONG,D. M. 1986. Oceanographer transform fault structure compared to that of surrounding oceanic crust - Results from seismic refraction data analysis. Journal of Geodynamics, 5, 79-102. BIBEE, L. D. & SHOR, G. G. 1976. Compressional wave anisotropy in the crust and upper mantle. Geophysical Research Letters, 3, 639-642. BOWN, J. W. & WHITE, R. S. 1994. Variation with spreading rate of oceanic crustal thickness and geochemistry. Earth and Planetary Science Letters, 121,435~449. &- 1995. Effect of finite extension rate on melt generation at rifted continental margins. Journal of Geophysical Research, 100, 18 01118 059. CANN, J. R. 1974. A model for oceanic crustal structure developed. Geophysical Journal of the Royal Astronomical Society, 39, 169-187. CANNAT,M. 1993. Emplacement of mantle rocks in the seafloor at mid-ocean ridges. Journal qf Geophysical Research, 98, 41634172, 1993. 1996. How thick is the magmatic crust at slow spreading oceanic ridges? Journal of Geophysical Research, 101, 2847-2857, 1996. CARLSON, R. L. & MILLER, D. J. 1997. A new assessment of the abundance of serpentinite in the oceanic crust. Geophysical Research Letters, 24, 457-460. CHIAN, D. & LOUDEN, K. E. 1994. The continentocean transition across the southwest Greenland margin. Journal of Geophysical Research, 99, 9117-9135. - & REID, I. 1995. Crustal structure of the Labrador Sea conjugate margins and implications for the formation of nonvolcanic continental margins. Journal of Geophysical Research, 100, 24 239-24 253. CHRISTENSEN, N. I. 1978. Ophiolites, seismic velocities

SERPENTINIZATION IN OCEANIC CRUST and oceanic crustal structure. Tectonophysics, 47, 131-167. CORMIER, M.-H., DETRICK, R. S. & PURDY, G. M. 1984. Anomalously thin crust in oceanic fracture zones: new seismic constraints from the Kane fracture zone. Journal of Geophysical Research, 89, 10249 10266. COULTON, A. J., HARPER, G. D. & O'HANLEY, D. S. 1995. Oceanic versus emplacement age serpentinization in the Josephine ophiolite: Implications for the nature of the Moho at intermediate and slow spreading ridges. Journal of Geophysical Research, 100, 22 245-22 260. DETRICK, R. S. & PURDY, G. M. 1980. The crustal structure of the Kane fracture zone from seismic refraction studies. Journal of Geophysical Research, 85, 3759-3778. , CORMIER, M.-H., PRINCE, R. A., FORSYTH, D. W. & AMBOS, E. L. 1982. Seismic constraints on the crustal structure of the Vema fracture zone. Journal of Geophysical Research, 87, 599-610. , WHITE, R. S. & PURDY, G. M. 1993. Crustal structure of North Atlantic fracture zones. Reviews of Geophysics, 31, 439-458. , COLLINS, J., STEPHEN, R. & SWIFT, S. 1994. In situ evidence for the nature of the seismic layer 2/3 boundary in oceanic crust. Nature, 370, 288-290. DICK, H. J. B., SCHOUTEN,H., MEYER,P. S., GALLO,D. G., BERGH, H., TYCE, R., PATRIAT, P., JOHNSON, K. T. M., SNOW, J. & FISHER, A. 1991a. Tectonic evolution of the Atlantis II Fracture Zone. In: VON HERZEN, R. P., ROBINSON, P. T. et al.,

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(Ocean Drilling Program), 359-398. MEYER,P. S., BLOOMER,S., KIRBY, S., STAKES, D. & MAWER, C. 1991b. Lithostratigraphic evolution of an in-situ section of oceanic layer 3. In." YON HERZEN, R. P., ROBINSON, P. T. et al.,

Proceedings of the Ocean Drilling Program, Scientific Results, 118, College Station, TX (Ocean Drilling Program), 439-538. DYMENT, J., ARKAN1-HAMED, J. & GHODS, A. 1997. Contribution of serpentinized ultramafics to marine magnetic anomalies at slow and intermediate spreading centres: insights from the shape of the anomalies. Geophysical Journal International 129, 691-701. EVANS, B. W., JOHANNES, W., OTERDOOM, W. H. & TROMMSDORF, V. 1976. Stability of chrysotile and antigorite in the serpentine multisystem. Schweit-

zerische Mineralogische Petrologische Mitteilungen, 56, 79-93. GREVEMEYER, I., WEIGEL, W., WHITMARSH, R. B., AVEDIK, F. & DEGHANI, G. A. 1997. The Aegir Rift: crustal structure of an extinct spreading axis. Marine Geophysical Researches, 19, 1-23. HENSTOCK, T. J., WHITE, R. S. & MCBRIDE, J. H. 1996. Along-axis variability in crustal accretion at the Mid-Atlantic Ridge: Results from the OCEAN study. Journal of Geophysical Research, 101, 13673 13688. HEss, H. H. 1962. History of ocean basins. In." ENGEL, A. E., JAMES, H. L. & LEONARD,B. F. Petrologic

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Studies, Burlington Volume. Geological Society of America, Boulder, Colorado, pp 599-620. HOREN, H., ZAMORA, M. & DUBUlSSON, G. 1996. Seismic waves velocities and anisotropy in serpentinized peridotites from Xigase ophiolite: Abundance of serpentine in slow spreading ridge. Geophysical Research Letters, 23, 9-12. HUANG, P. Y. & SOLOMON,S. C. 1988. Centroid depths of mid-ocean ridge earthquakes: dependence on spreading rate. Journal of Geophysical Research, 93, 13445-13477. KEMPNER, W. C. & GETTRUST, J. F. 1982. Ophiolites, synthetic seismograms and oceanic crustal structure. Journal of Geophysical Research, 87, 84478476. KERN, H., LIu, B. & PoPP, T. 1997. Relationship between anisotropy of P and S wave velocities and anisotropy of attenuation in serpentinite and amphibolite. Journal of Geophysical Research, 102, 3051-3065. KLEIN, E. M. & LANGMUIR, C. H. 1987. Global correlations of ocean ridge basalt geochemistry with axial depth and crustal thickness. Journal of Geophysical Research, 92, 8089-8115. LANGMUIR, C. H., KLEIN, E. M. & PLANK, T. 1992. Petrological systematics of mid-ocean ridge basalts: Constraints on melt generation beneath ocean ridges. In: PHIPPS MORGAN, J., BLACKMAN, D. K. & SINTON,J. M. (eds) Mantle Flow and Melt Generation at Mid-Ocean Ridges. Geophysical Monograph 71, American Geophysical Union, Washington, DC, 183-280. LOUDEN, K. E., OSLER, J. C., SRIVASTAVA, S. P. & KEEN, C. E. 1996. Formation of oceanic crust at slow spreading rates: New constraints from an extinct spreading center in the Labrador Sea. Geology, 24, 771-774. LUDWIG, W. J. & RABINOWITZ,P. D. 1980. Structure of the Vema fracture zone. Marine Geology, 35, 99110. MACDONALD, A. W. & FYFE, W. S. 1985. Rate of serpentinization in seafloor environments. Tectonophysics, 116, 123-135. MCKENZIE, D. & O'NloNs, R. K. 1991. Partial melt distributions from inversion of rare earth element concentrations. Journal of Petrology, 32, 10211091. MILLER, D. J. & CHRISTENSEN, N. I. 1997. Seismic velocities of lower crustal and upper mantle rocks from the slow-spreading Mid-Atlantic Ridge, south of the Kane Transform (MARK). In: KARSON, J. A., CANNAT, M., MILLER, D. J. & ELTHON, D. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 153, College Station, TX (Ocean Drilling Program), 437-454. MINSHULL, T. A., WHITE, R. S., MUTTER, J. C., BUHL, P., DETRICK, R. S., WILLIAMS,C. A. & MORRIS, E. 1991. Crustal structure at the Blake Spur fracture zone from expanding spread profiles. Journal o.1" Geophysical Research, 96, 9955-9984. MULLER, M. R., ROBINSON, C. J., MINSHULL, T. A., WHITE, R. S. & BICKLE, M. J. 1997. Thin crust beneath ocean drilling program borehole 735B at the Southwest Indian Ridge? Earth and Planetary

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Science Letters, 148, 93-107. NICOLAS, A. 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Academic Publishers, Dordrecht. NIU, Y. & HEKINIAN, R. 1997. Spreading-rate dependence of the extent of mantle melting beneath mid-ocean ridges. Nature, 385, 326-329. O'HANLEY, D. S. 1996. Serpentinites: Records of Tectonic' and Petrological History. Oxford University Press, New York. OSLER, J. C. & LOUDEN, K. E. 1992. Crustal structure of an extinct rift axis in the Labrador Sea: preliminary results from a seismic refraction survey. Earth and Planetary Science Letters, 108, 243-258. PALMER, M. R. 1996. Hydration and uplift of the oceanic crust on the Mid-Atlantic Ridge associated with hydrothermal activity: Evidence from boron isotopes. Geophysical Research Letters, 23, 3479-3482. Penrose Conference on Ophioloites. 1972. Geotimes, 17, 24-25. PINHEIRO, L. M., WHITMARSH, R. B. & MILES, P. R. 1992. The ocean-continent boundary off the western continental margin of Iberia-II. Crustal structure in the Tagus Abyssal Plain. Geophysical Journal International, 109, 106-124. POTTS, C. G., WHITE, R. S. & LOUDEN, K. E. 1986a. Crustal structure of Atlantic fracture zones, II. The Vema fracture zone and transverse ridge. Geophysical Journal of the Royal Astronomical Society, 86, 491-513. - - , CALVERT,A. J. & WHITE, R. S. 1986b. Crustal structure of Atlantic fracture zones, III. The Tydeman fracture zone. Geophysical Journal of the Royal Astronomical Society, 86, 909-942. RAirr, R. W. 1963. The crustal rocks. In: HILL, M. N. (ed.) The Sea, Vol. 3, Wiley Interscience, New York, 85-102. REID, I. 1994. Crustal structure of a non-volcanic rifted margin east of Newfoundland. Journal of Geophysical Research, 99, 15 161-15180. - - , & JACKSON, H. R. 1981. Oceanic spreading rate and crustal structure. Marine Geophysical Researches, 5, 165-172. ROBINSON, C. J., WHITE, R. S., BICKLE, M. J. & MINSHULL, T. A. 1996. Restricted melting under the very slow-spreading Southwest Indian Ridge. In: MACLEOD, C. J., TYLER, P. A. & WALKER, C. L. (eds) Tectonic, Magmatic, Hydrothermal and

Biological Segmentation of Mid-Ocean Ridges. Geological Society, London, Special Publications, 118, 131 141. RONA, P. A., BOUGAULT,H., CHARLOU,J. L., APPRIOU,

P., NELSON, T. A., TREFRY, J. H., EBERHART, G. L., BARONE, A. & NEEDHAM, H. D. 1992. Hydrothermal circulation, serpentinization, and degassing at a rift valley-fracture zone intersection: Mid-Atlantic Ridge near 15°N, 45°W. Geology, 20, 783-786. , WIDENFALK, L. • BOSTR6M, K. 1987. Serpentinized ultramafics and hydrothermal activity at the Mid-Atlantic Ridge crest near 15°N. Journal of Geophysical Research, 92, 1417-1427. SINHA, M. C. & LOUDEN, K. E. 1986. The Oceanographer fracture zone, I. Crustal structure from seismic refraction studies. Geophysical Journal of the Royal Astronomical Society, 75, 713-736. VANKO, D. A. & STAKES,D. S. 1991. Fluids in oceanic layer 3: evidence from veined rocks, hole 735B, Southwest Indian Ridge. In: VON HERZEN, R. P., ROBINSON, P. T. et al. Proceedings of the Ocean Drilling Program, Scientific Results, 118, College Station, TX (Ocean Drilling Program), 181-215. WHITE, R. S., MCKENZIE, D. & O'NIONS, R. K. 1992. Oceanic crustal thickness from seismic measurements and rare earth element inversions. Journal of Geophysical Research, 97, 19683 19715. WHITMARSH, R. B. & CALVERT, A. J. 1986. Crustal structure of Atlantic fracture zones-I. The Charlie-Gibbs fracture zone. Geophysical Journal of the Royal Astronomical Society, 85, 107-138. - & SAWYER, D. S. 1996. The ocean/continent transition beneath the Iberia abyssal plain and continental-rifting to seafloor-spreading processes. In." WHITMARSH, R. B., SAWYER, O. S., KLAUS, A. & MASSON, D. G. (eds) Proceedings of

the Ocean Drilling Program, Scientific Results, 149, College Station, TX (Ocean Drilling Program), 713-733. - - , MILES, P. R. & MAUFFRET, A. 1990. The ocean-continent boundary off the western continental margin of Iberia-I Crustal structure at 40°30'N. Geophysical Journal International, 103, 509-531. - - , WHITE, R. S., HORSEFIELD, S. J., SIBUET,J.-C., RECQ, M. & LOUVEL, V. 1996. The oceancontinent boundary off the western continental margin of Iberia: Crustal structure west of Galicia Bank. Journal of Geophysical Research, 101, 28 291-28 314. WOLFE, C. J., PURDY, G. M., TOOMEY, D. R. & SOLOMON,S. C. 1995. Microearthquake characteristics and crustal velocity structure at 29°N on the Mid-Atlantic Ridge: The architecture of a slow spreading segment. Journal of Geophysical Research, 100, 24 449-24 472.

Tracing the evolution of hydrothermal fluids in the upper oceanic crust: Sr-isotopic constraints from D S D P ] O D P Holes 504B and 896A D. A. H . T E A G L E , J. C. A L T & A. N. H A L L I D A Y

Department o f Geological Sciences, 2534 C.C. Little Building, The University o f Michigan, Ann Arbor, M I 48109-1063, USA

Abstract: Whole rock and secondary mineral 87Sr/86Sr determinations from ODP Holes 504B and 896A have been compiled into a profile for the upper 2 km of in situ ocean crust to constrain the composition and evolution of recharge and discharge fluids. There is a gradual trend from elevated 87Sr/S6Sr whole rock ratios at the top of the crust towards primary MORB compositions at the base of the sheeted dykes. Rocks from the mineralized lavadyke transition zone have 87Sr/86Sr ratios clustering in the range of black smoker fluids. Igneous and alteration phases are not in Sr-isotopic equilibrium, therefore whole rock strontium isotopic ratios reflect the proportion of secondary minerals and the extent of plagioclase recrystallization. The 87Sr/S6Sr of anhydrite directly records the chemical evolution of seawater during recharge. There is a slight decrease in anhydrite SVSr/86Srwith depth in the volcanic rocks but anhydrite from the lowermost dykes yields diabasic 87Sr/86Sr ratios. Epidotes in cross-cutting veins have higher S7Sr/86Sr than the background alteration recorded by amphibole in the lower dykes. This elevated signature reflects the 87Sr/S6Sr of 504B black smoker-type fluids and requires less fluid evolution than the background alteration, suggesting channelling of recharge fluids locally. The composition of seawater is in part, controlled by hydrothermal interactions with the ocean floor, yet the detailed mechanisms and transport paths underlying these alteration processes are not well understood. The different and well characterised strontium isotopic compositions of ocean ridge basalts (87Sr/ 86Sr~0.7025) and seawater (87Sr/ 86Sr~0.7092), and the relatively similar strontium concentrations of these reservoirs ([Sr]MORB~90 ppm; [Sr]Seawater~8 ppm), makes Sr an ideal tracer of seawater-basalt interaction and hydrothermal alteration (Spooner et al. 1977; Bickle & Teagle 1992). The analysis of whole rock and mineral samples for 87 Sr/ 86 Sr can yield constraints on the co-existing fluid compositions and records the evolving conditions of fluid-rock interaction. Alteration within ocean floor rocks is manifest by two basic phenomena: (1) the replacement of the constituent primary igneous phases by secondary minerals and (2) the precipitation of hydrothermal minerals in vugs and cross-cutting veins. As different secondary minerals form under contrasting physical and chemical conditions, their presence and mode of occurrence in ocean floor rocks provide information on the local environment at the time of their formation. Here we present a compilation of 87Sr/86Sr determinations for whole rock samples and

mineral separates from ODP Holes 504B and 896A that defines a strontium isotopic profile for the upper 2 km of in situ hydrothermally altered ocean crust, and constrains the composition and evolution of basement fluids.

Geological Setting of DSDP]ODP Holes 504B and 896A D S D P / O D P Hole 504B is located in 5.9Ma crust, 200 km south of the intermediate spreading rate Costa Rica Rift in the eastern equatorial Pacific (Fig. l). It is the deepest hole in the oceanic crust and penetrates more than 1800 m into basement t h r o u g h the entire volcanic sequence, and nearly to the base of the sheeted dyke complex (Alt et al. 1993). Hole 504B provides a reference section for the uppermost ocean crust, and detailed investigations have led to a general model for the h y d r o t h e r m a l alteration (Alt et al. 1986a; Alt 1995; A l t e t al. 1996a). Hole 504B penetrates 274.5 m of sediments, a 571.5 m volcanic section comprising pillow lavas and massive flows, breccias, and possible local dykes or sills in the lower half of the section; a 209 m transition zone of dykes and lava screens; and 1056 m of a sheeted dyke complex. Core recovery averages 30% in the volcanic section, 25% through the transition zone and 14% in the sheeted dyke complex.

TEAGLE,D. A. H., ALT, J. C. & HALLIDAY,A. N. 1998. Tracing the evolution of hydrothermal fluids in the upper oceanic crust: Sr-isotopic constraints from DSDP/ODP Holes 504B and 896A. In: MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 81-97

81

82

D.A. H. TEAGLE E T AL.

10°N

trated by Hole 896A. The down-hole logging interpretations are not without ambiguities however.

_

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Site 896 is approximately 1 km to the southeast of Hole 504B (Alt et al. 1993) and this site was previously occupied for sediment pore-water studies during Leg 111 (Site 678; Becker & Sakai 1988). Hole 896A is situated on a bathymetric high that coincides with the topographic high point of a tilted basement fault block and a local heat-flow maximum (Fig. 2; Langseth et al. 1988). Sediment pore water profiles show that basement fluids well-up at high heat-flow sites (e.g. Site 678/896A) and that bottom seawater is being drawn into the crust through the sediment blanket in zones of low heat-flow (Mottl 1989). Hole 896A penetrated to 469 m below seafloor (mbsf), through 179 m of sediment and 290 m of basement, with an average basement recovery of 28% (Alt et al. 1993). The basaltic basement sampled comprises interlayered pillow lavas (57%), massive flows (38%), breccias (5%), and two sub-vertical dykes, a similar distribution to materials recovered from the uppermost crust in Hole 504B (see Alt et al. 1996c). Breccias comprise 5% of the core recovered from Hole 896A; however, interpretation of Formation microscanner (FMS) and Geochemical Logging Tool (GLT) data recovered on Leg 148 (Brewer et al. 1995) suggests that breccias may make up almost 50% of the volcanic stratigraphy pene-

The volcanic section can be divided into upper and lower alteration zones but the basalts throughout the extrusive sequence are generally only slightly altered (5-15% secondary minerals), except for hyaloclastic and tectonic breccias which are strongly recrystallized and comprise ~13% of the recovered material in Hole 504B (Alt et al. 1996c). The upper half of the volcanic section exhibits broadly similar features at both Sites 504 and 896 (Alt et al. 1996c; Teagle et al. 1996). This portion of the basement was open to the seafloor and the basalts interacted with large volumes of oxidizing ocean water at low temperatures (< 100°C) resulting in the development of iron-oxyhydroxide-rich and celadonitic halos around veins (see Alt et al. 1996c). Saponite (Mg-smectite) and calcium carbonate are the most abundant secondary minerals (Fig. 3) however, and are present in veins, as breccia cements and filling pore space. Lavas deeper in the volcanic section were altered under more reducing conditions with more restricted fluid flow and lower fluid fluxes. Alteration of the volcanic section probably began at the ridge axis, but continued for at least several million years (Mottl & Wheat 1994; Alt et al. 1996c). Hydrothermal circulation continues in the uppermost hundred metres of the basement today, and cementation and alteration of this portion of the crust may be occurring presently (Fisher et al. 1990, 1994; Teagle et al. 1996). The lithologic transition zone marks the gradational contact between volcanic flows and 100% dykes. The rocks of the transition zone are mainly highly fractured and brecciated hydrothermally altered lavas and dykes. The physical properties of the crust measured in situ change significantly across the transition zone, with sonic velocities increasing sharply, whereas bulk permeability and porosity drop by orders of magnitude (Becker et al. 1989). A mineralized stockwork-like zone containing abundant veins of quartz and sulphide minerals occurs in a highly fractured and brecciated pillow unit at 910-928 mbsf, within the transition zone. These rocks were altered in a sub-surface mixing zone, where 250-350°C hydrothermal fluids upwelling through the dykes mixed with cooler seawater solutions circulating in the overlying volcanic section. Rocks of the upper dyke section are light to dark grey and are variably recrystallized

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(10-100%) to sub-greenschist facies assemblages. Fine to medium grained diabase comprises the lower sheeted dykes and all rocks exhibit some effects of hydrothermal alteration at higher temperatures, approaching amphibolite facies conditions. The least altered rocks are dark grey and affected by a pervasive 'background' alteration, and are 10-40% recrystallized to greenschist-facies secondary minerals. Light coloured alteration halos (1-20mm) are common around amphibole + chlorite veins and are more intensely recrystallized (30-100 %) than the adjacent dark grey host diabase. Marked depletions in base metals (Cu, Zn) and sulphur contents in the vein halos, the presence of anhydrite filling pore-space in the rocks, and the development of secondary magnesio-hornblende, calcic-plagioclase and clinopyroxene, indicate that the rocks from the lower sheeted dyke complex form part of the 'reaction zone' where hydrothermal fluids acquire their chemi-

cal signatures through fluid-rock interaction at temperatures > 350°C (Alt et al. 1996b; Vanko et al. 1996). The alteration of the ocean crust comprises a continuum of overprinting thermal and fluid flow processes as the crust matures away from the zone of active intrusion and volcanism. The sequence of hydrothermal alteration in the transition zone and the dyke section can be summarized in five stages (Alt et al. 1986a; Laverne et al. 1995; Alt et al. 1996b). Secondary calcic plagioclase (labradorite-anorthite) and magnesio-hornblende formed in patches and vein halos in the lower dykes at temperatures > 400°C followed by the precipitation of chlorite and actinolitic amphibole in veins and the development of greenschist minerals (actinolitic amphibole, albite-oligoclase, chlorite, talc, and titanite) in the groundmass during axial hydrothermal alteration at temperatures of ~250350°C. These assemblages are cross-cut by

84

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Fig. 3. Lithostratigraphy of Hole 504B, showing the distribution of secondary minerals with depth (from Alt et al. 1996a). Hole 896A penetrates 290m sub-basement and has a similar distribution of secondary minerals to the upper part of Hole 504B. UPA- Upper Pillow Alteration; LPA- Lower Pillow Alteration. Heavy horizontal line at ~640 m sub-basement shows the location of the mineralized stockwork within the transition zone.

quartz, epidote and sulphides veins, and the sulphide mineralization that formed the stockwork in the transition zone. Later, as the crust moved off-axis into a recharge zone, penetration of seawater into still warm rock resulted in the precipitation of anhydrite in cracks and within the sheeted dykes, and the local replacement of plagioclase by anhydrite. Finally uncommon zeolites formed in fractures within the volcanic rocks during later off-axis alteration at lower temperatures (< 250°C).

A strontium isotopic profile through the upper ocean crust A profile of previously published and new whole rock 87Sr/86Sr ratios through the upper ocean crust penetrated by Holes 504B and 896A is shown in Fig. 4. Strongly brecciated rocks with either a hyaloclastic, tectonic or hydrothermal origin are also displayed. The strontium isotopic composition of fresh MORB (~0.7025), and 5.9 Ma and present-day seawater, of 0.70893 and

0.70917, respectively (Hodell et al. 1991) are plotted for comparison. Within the volcanic section of Hole 504B there is a gradual trend in the whole rock samples from moderately elevated ratios (87 Sr/ 86 Sr,.~0.7035-0.7055) at the top of the crust toward primary MORB compositions (87Sr/86Sr~0.7025) at the base of the volcanic pile. This would suggest a decreasing seawater effect with increasing depth in the lavas, consistent with other chemical, mineralogical, and isotopic evidence (Alt et al. 1986a,b, 1989). There is a sampling bias in the Hole 504B analyses towards the least altered rock-types and data from that hole yield 87 Sr/ 86 Sr ratios significantly lower than altered samples from Hole 896A. The enrichment in radiogenic 878r and the range displayed by the different alteration types decreases with depth. Below 250m, sub-basement whole rock samples from the upper pillows have only slightly elevated STSr/S6Sr ratios (<0.7035) and there are only small variations between the different alteration types. Higher in the crust, juxtaposed subsam-

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Fig. 4. 87Sr/86Sr profile of the upper oceanic crust compiled from whole rock data from Holes 504B (open circles) and 896A (solid squares); Hydrothermal, hyaloclastic and tectonic breccias are shown by stars (Barrett & Friedrichsen 1982; Barrett 1983; Hart & Mottl 1983; Friedrichsen 1985; Staudigel & Hart 1985; Kawahata et al. 1987; Kusakabe et al. 1989; Shimizu et al. 1989; Alt et al. 1996a,b; Teagle et al. 1996 and Teagle, unpublished data). The shaded region in the volcanic sequences shows the range of Sr-isotopic compositions yielded by composite samples from Sites 417 and 418 in 120 Ma-old Atlantic crust (Staudigel et al. 1995). Depths have been normalized to match the ~580 m pillow lava section present in Hole 504B. pies with contrasting alteration assemblages (e.g. Mg-saponite, celadonite, calcium carbonate) have quite different strontium isotopic compositions though a consistent relationship of elevated 87Sr/86Sr and a particular type of alteration is not observed (Teagle e t al. 1996). Brecciated samples from both Holes 504B and 896A have moderately to highly elevated 87Sr/86Sr ratios, indicating substantial incorporation of seawater strontium (65-95%). A highly altered hyaloclastic breccia from the u~permost pillow lavas of Hole 504B has a r/S6Sr similar to clay and carbonate mineral separates, indicating the complete recrystallization to secondary mineral phases through interaction with basement fluids having near-

seawater strontium isotopic compositions. Presently there is only limited data for altered rocks from the lower lava sequences in Hole 504B, though a partially altered hyaloclastic breccia from ~500 m sub-basement has a moderately elevated 87Sr/86Sr ratio (~0.7042) indicating that recharge fluids retain a significant seawater-St component down to the base of the lava pile. Within the transition zone and associated stockwork mineralisation all samples have elevated 87Sr/86Sr ratios (~0.703-0.707) recording a significant component of seawater-derived strontium in these rocks. Data are generally clustered in the range 8 7 Sr/ 8 6 S r = 0.7030-0.7045, similar to the variation displayed by high

86

D . A . H . TEAGLE ET AL.

temperature black smoker vent fluids (Palmer & Edmond 1989). The transition zone represents a region of fluid mixing and some of the highest 87Sr/86Sr ratios presumably reflect the dilution of u~welling hydrothermal fluids having low r/S6Sr ratios with seawater. There is significant scatter in 87Sr/86Sr (~0.7025-0.7055) of the uppermost sheeted dykes (~790 1320m sub-basement), though the range in 87Sr/S6Sr decreases with depth. Less altered, grey background diabase are generally only little altered from MORB isotopic compositions, but breccias and localised highly altered atches and vein halos generally fall in the range Sr/S6Sr ~0.703-0.704. The same effect of decreasing seawater influence is seen in other chemical and isotopic tracers of hydrothermal interaction (e.g., ~SLSO,634S; see Alt et al. 1986b, 1989). The lower dykes (1325-1835 m sub-basement) have a restricted range of whole-rock 87Sr/S6Sr ratios of 0.70265-0.70304. These are only slightly elevated toward seawater composition relative to the lowest value of 0.7025 reported from Hole 504B, which is presumably near the primary MORB value for Site 504 (Shimizu et al. 1989). There is a slight trend of an increasing alteration baseline, with no diabase from the comprehensively sampled section drilled on Leg 148 (~1735-1835 m sub-basement) yielding 87Sr/S6Sr <0.70266, whereas such values are common higher in the dyke section (Alt et al. 1996b). This slight elevation in the 878r/S6Sr of the least altered rocks, must reflect either an increase in the proportion of Sr-bearing secondary minerals or a more radiogenic Sr-isotopic composition of the reacting hydrothermal fluids. Because the rocks are not completely recrystallized and primary minerals retain their igneous Sr isotopic compositions, there must be significant Sr-isotopic disequilibrium within the diabase, indicating that alteration and strontium exchange were kinetically retarded even at the deepest levels of the sheeted dykes. Strontium

isotopic

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temperature secondary minerals

The Sr-isotopic composition of carbonate, clay minerals, and zeolites from the uppermost ocean crust (< 700m sub-basement) at Sites 504 and 896 are shown in Fig. 5. Mg-saponite is the most abundant secondary mineral in Holes 504B and 896A and occurs replacing mesostasis, as a breccia cement and most commonly filling fractures and vugs. Although individual rocks may not show any significant gain of Mg, the

abundance of saponite veins and the strong Mgenrichment of breccias indicate that the volcanic section as a whole has taken up Mg during saponite formation. Most saponites from both holes have high, near-seawater S7Sr/S6Sr (~0.7084-0.7091) but, some of these samples, have high Rb contents (> 7ppm) and high 87Rb/s6Sr (~10) for which a significant age correction may be necessary depending on the time of formation of the clay. Insufficient data are presently available to construct a rigorous isochron but the clays can be no more than 5.9 Ma old. Applying this age correction these saponites could have SVSr/S6Sr initial ratios as low as 0.70825. A completely recrystallized hyaloclastic breccia from the top of the ocean crust has a Sr-isotopic composition indistinguishable from clay minerals occurring in veins. Clay minerals that have 87Sr/S6Sr ratios much lower than 0.7082 are either from rare fine veins of fibrous saponite or from one sample that contains a significant proportion of partially altered glass which retains its primary basaltic Sr. The veins of fibrous saponite yield ratios similar to the ratios in the host rocks (Teagle et al. 1996) indicating that fluids may react locally to rock-dominated compositions, perhaps in regions of much more restricted fluid access. Calcite, and aragonite in the upper pillows are common secondary minerals filling vugs, veins and as breccia cements in the volcanic section (Fig. 3). Uncommon calcite is also present in the hydrothermally mineralized Transition Zone (Alt et al. 1986a; Hart et al. 1994). Carbonate samples from Hole 896A and the upper part of Hole 504B have remarkably similar Sr isotopic compositions (87Sr/86Sr= 0.7084-0.7088), which are apparently insensitive to depth of occurrence, carbonate mineralogy, habit or oxygen isotopic composition (Teagle et al. 1996). Two samples of aragonite from uncommon fine veinlets in Hole 896A yield significantly less radiogenic Sr ratios than the other carbonates ( 87 Sr/,86 Sr=0.707943.7084). The strontium isotopic composition of seawater has increased monotonically for the last 45 Ma (Burke et al. 1982; Hodell et al. 1991) and present-day and 5.9 Ma seawater have 87 Sr/ 86 Sr ratios of 0.70919 and 0.70895, respectively (Hodell et al. 1991). All carbonate samples from Holes 504B and 896A have 87Sr/86Sr lower than that of 5.9Ma seawater (Fig. 5), though significantly higher than fresh MORB or altered basalt ( 87 Sr/ 86 Sr~0.70250.7055). Carbonate 87Sr/86Sr ratios lower than seawater values require the incorporation of basaltic strontium ( 87 Sr/ 86 Sr=0.7025-0.7027) and preclude the use of the seawater strontium curve for constraining

TRACING HYDROTHERMAL FLUID EVOLUTION

87

8 r/86Sr 0.703

0.705 ,

I

,

0.707 , I

0.709 I

Ii II 1 O0 t'ID

°D

iii I I I

E 200 or) ..13

_& .m

300 -

[]

or) or) I11 Ik-ii1 4 0 0 E v L o w e r Pillow Alteration

t-o.. 5 0 0 (D £3

o'1 " o

600=~ ~,~:.~NNg~ ;~ ~ , ~ ~

Tra n s itio n Zone

II II

700

Fig. 5. 87Sr/86Sr of secondary minerals from the uppermost ocean crust recovered in Holes 504B (open symbols) and 896A (filled symbols). Shaded area delineates the range of whole rock 878r/86Sr from these holes. CirclesMg-saponite: squares-calcium carbonate; diamonds-zeolites; stars breccias. Mineral data from Hart et al. (1994), Teagle et al. (1996) and Teagle unpublished data.

the time of deposition of these carbonates. Pore waters in the overlying sediments may have 87Sr/*6Sr depleted with respect to modern seawater, but the lower value of this depletion is limited by the age of the oldest sediment (< 5.9 Ma), and hence reactions in the sediment column cannot be the cause of the depleted radiogenic strontium ratios in the carbonate veins. Depending on the time of carbonate precipitation (0-5.9 Ma B.P.) and the nature of the host rock with which basement fluids interacted (STSr/86Sr=0.7027-0.7055), approximately 210% of the strontium in the carbonates must have been derived from the basalt. The limited range of 87Sr/86Sr ratios recorded from the majority of carbonates indicate that basement fluids, at the time of carbonate deposition, were buffered to this strontium isotopic composition by wall rock interaction, at least in the upper few hundred meters of the ocean crust. For the aragonite with the lowest 87Sr/86Sr (~0.7079), the proportion of rock-derived strontium is 15-

30%, approximately twice that in the other carbonates. Calcites from the t r a n s i t i o n zone have 87Sr/S6Sr ratios (~0.7041-0.7048) intermediate to seawater and either primary MORB or Hole 504B black smoker fluids (~0.7034-0.7038; see later). The ~13C of calcites from this portion of the crust (4-11%o PDB; (Alt et al. 1985a) are similar to values for CO2 in black smoker fluids and indicate a mantle CO2 source as opposed to seawater (Welhan & Craig 1979; Lilley et al. 1993). 5180 and fluid inclusion measurements indicate elevated temperatures of formation for these calcites (T~180-200°C; Alt et al. 1986b) compared with carbonates shallower in the ocean crust. Isotopic and petrographic evidence indicate that this section of the crust is a zone of mixing between hot upwelling hydrothermal fluids and recharge seawater, and Sr-isotopic ratios indicate that these calcites formed by the mixing of upwelling hydrothermal fluids with a minor component of seawater. A large (~5cm) vug filled with euhedral

88

D. A. H. TEAGLE E T AL.

analcite and fibroradial natrolite that grew within a fluid-filled cavity in the pillow lavas was recovered from ~250m sub-basement in Hole 896A (Teagle et al. 1996). The natrolite has a 87Sr/86Sr measured ratio of 0.707440, intermediate between MORB and seawater, but the a n a l c i t e r a t i o is e x t r e m e l y r a d i o g e n i c (S7Sr/S6Sr=0.70914) close to the Sr isotopic composition of modern seawater. The natrolite has a low STRb/86Sr (~0.02) and there is no significant age correction but the analcite has a moderately high S7Rb/s6Sr (~3.62) and could have an initial ratios as low as 0.7088 (for t = 5.9 Ma). These two minerals do not form a geologically sensible isochron (i.e. requires t > 5.9 Ma) and the near seawater-like STSr/S6Sr ratio of the analcite suggests that this mineral may have freely exchanged Sr with seawater up to the present day. Carbonate is commonly associated with the formation of saponite in veins of Holes 504B and 896A, but also postdates saponite in some veins. The similarity in 87Sr/86Sr of most saponites and carbonates indicates that both minerals formed from fluids that had undergone a similar extent of water-rock interaction. In Hole 896A oxygen isotopic evidence indicates that carbonates formed under two differing temperature regimes, at approximately 25-35°C and 45-65°C (Teagle et al. 1996). There is no direct constraint on relative or absolute timing of the carbonates formed at the different temperatures, but the general evolution of ridge-flank circulation from early low temperatures to later higher temperatures suggests that the lower-temperature vein carbonates formed before the higher temperature carbonates. Temperatures of formation estimated for the higher temperature carbonates in Hole 896A are similar to the present-day temperatures in the upper crust (Alt et al. 1993) suggesting that these carbonates could have formed under the current thermal conditions, which result from the location of Site 896 in a zone of ridge-flank hydrothermal upwelling. The shaded region on Fig. 4 shows the range of 87Sr/S6Sr compositions for composite samples of volcanic rocks from Holes 417A, 417D, and 418A in 120 Ma crust from the western Atlantic Ocean (Staudigel et al. 1995). Depths have been normalized to coincide with the ~580 m volcanic sequence penetrated in Hole 504B. The 417/418 analyses have much larger proportions of seawater-Sr than rocks from Hole 504B or 896A. Adjustment of the Hole 504B/896A data to 120 Ma results in only minimal overlap due to their low Rb/Sr ratios (see Teagle et al. 1996). Only brecciated rocks or carbonate-bearing basalts

from these sites coincide with the Atlantic data. This discrepancy is partially due to the different modes of sampling (whole rock + mineral separates v. mixtures of powders) from these different studies however, the 417/418 composites are strongly biased by extremely radiogenic samples recovered from the highly altered Hole 417A. This hole was drilled into an exposed abyssal high, and represents an extreme style of highly oxidative alteration at high water/rock ratios (Alt & Honnorez 1984; Staudigel et al. 1995). If only data from Holes 417D and 418A, which display assemblages more typical of low temperature ocean floor alteration, are considered there are only slight differences between the 417/418 and 504B/896A datasets. The high rates of recovery from the Atlantic holes are indicative of strongly altered basalts, in which clay minerals thoroughly cement fractures and breccias, reducing the extent of disturbance during drilling. The anomaly between the measured heat flow and that calculated from slab-cooling models (Stein & Stein 1994) extends on average, out to crust of 60 Ma, indicating that heat removal by fluid convection continues to be a significant process for at least this time period. As heat transport is unlikely to occur without any fluidrock interaction and some mineral precipitation, it is entirely consistent that basalts from Sites 417/418 (~ 120 Ma) show slightly greater extents of alteration than rocks from Holes 504B and 896A (~5.9 Ma). As water/rock ratios are high and there are only small amounts of reaction, changes in the Sr-isotopic compositions of ridge-flank fluids may be difficult to detect, but these flank reactions could still have a significant effect on ocean chemistry because of the huge fluid fluxes involved (Mottl & Wheat 1994).

Alteration in the sheeted dykes and controls on the composition of hydrothermal fluids Isotopic and other differences between background altered and more highly altered diabase from the upper and lower sheeted dykes are investigated in Fig. 6. There is a general trend of increasing 87Sr/86Sr with alteration, as represented by the proportion of secondary minerals (Fig. 6a). Background dark grey diabases from the upper sheeted dykes tend to have lower proportions of secondary minerals than similar rocks from deeper in the dykes but more highly altered rocks from the upper dykes generally have higher 87sr/g6sr than rocks recrystallized to a similar extent from the lower dykes. The Sr-

TRACING HYDROTHERMAL FLUID EVOLUTION

A

A

~°°!

,

I

, A

I

, :~'~-'EC,

--'-'- 8 0

E

,

D

/

[]

/ i i: ;il'

v

{'-

I

=

89

I'~l

'

/'i /,'

//'

=

i , /" ~ ' " i r a

[]

6O

o

•

i [B

(1~ 40 : :':']~]"=]";:' ' "'~" ' - •" ..... ~. ~ ~~ u' ~~So

(1)

' ~ 2o '

B

100

/

'

" / ,• "I • o

I

o [] o •

'

I

0. 7 0 2 6

'

I

upper upper lower lower '

0.7030

dyke-dark d y k e - light dyke - dark d y k e - light

I

'

breccia A actinolite i C chlorite E epidote

I

I

0.7034

'0.703; I

L

Ii /

D

80 •

E ¢.~

60

//

o•~ o(~, • m• ~

x._ 4 0

" -

'"

niP

•

c~-_O__..o

• o

Or)

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2O

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0.7026

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,

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0.7034

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t

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0.7038

i

87Sr/86Sr Fig. 6. (A) 8 7 S r / 8 6 S r v. alteration (proportion of secondary minerals) for whole rock samples from the upper and lower dYkes from Hole 504B. The 87Sr/86Sr of actinolite (A), chlorite (C) and epidote (E) from hydrothermal veins are also shown. (B) 87Sr/S6Sr v. Sr concentration of diabase dikes from Hole 504B. Tie lines join neighbouring background dark grey diabase and more highly altered (light) samples. Data from Kawahata et al. (1987), Kusakabe et al. (1989), Alt et al. (1996a,b) and Teagle unpublished data. isotopic compositions of minerals separated from hydrothermal veins and a completely recrystallized hydrothermal breccia are also shown in Fig. 6a. Most tie lines between background and altered rocks trend towards the range of 87Sr/S6Sr delineated by the vein minerals, indicating that these manerals record the isotopic composition of the deep basement fluids (S7Sr/S6Sr~0.7027-0.7038). Samples from the lower dykes generally, though not exclusively, trend towards the limited range of actinolite S7Sr/~6Sr ratios (~0.7027-0.7032). The correlation between STSr/S6Sr and the abundance of secondary minerals indicates that significant Sr-isotopic disequilibrium existed between the bulk rocks and the hydrothermal fluid. The rocks are mixtures of secondary minerals, which record the isotopic composition of the hydrothermal fluid, with igneous phases that retain their original MORB isotopic composition. Analyses of mineral separates by Barrett & Friedrichsen (1987) show evidence for extreme Sr-isotopic disequilibrium in samples from the volcanic section of Hole 504B. The retention of MORB Sr-isotopic compositions by the primary igneous minerals implies that Sr

does not exchange diffusively with fluids over the time scale of hydrothermal interaction and that Sr is only exchanged by dissolution-reprecipitation mechanisms during the recrystallization to secondary mineral phases. There is no simple relationship between Sr v. STSr/S6Sr (Fig. 6b). Altered rocks from the lower dykes generally have Sr contents similar to the background diabase indicating that Sr is exchanged between the rocks and hydrothermal fluids with little change in the Sr concentration of either the rocks or fluids. However, some altered rocks are strongly enriched or depleted in Sr, indicating the dissolution or growth of Srbearing phases. The dissolution of Sr-bearing phases and the net loss of Sr from the whole rocks through alteration is more common in samples from the upper dykes than in diabase from the lowermost dykes. The localized modal changes in the alteration halo around a thin actinolite vein are shown in Fig. 7 and compared with variations in Sr content and isotopic composition. The most prominent difference between the vein halo and the surrounding background diabase is the large increase in secondary minerals, particularly the

90

D.A. H. TEAGLE E T AL.

90

z~

E Q.. v 70 I

<350 -~

go /4

.

0.7032

e13 ..~

"~

130 I 0

0.7030

o 10 5

_

°4/°o4 []

r

0.7028

o t4~'#l~g

eo .............. o O~"13~'g

__<#

Background dark gray diabase

i

~ i

Vein halo i 1 cm

Ioo (/)-.4

O. 7026

"

0.7024

Amphibole vein i

Fig. 7. Sketch of Sample 239R-1, 45-51 cm from Hole 504B from which the dark grey background diabase, vein halo, and amphibole vein were separated and analysed. Amphibole, and secondary sodic and calcic plagioclase increase in the alteration halo, whereas chlorite decreases relative to the adjacent dark grey rock. Modal proportions shown by open circles. Whole rock and amphibole vein Sr concentrations and 87Sr/86Sr indicated by open triangles and squares, respectively.

proportion of actinolite replacing clinopyroxene and interstitial material. The proportion of secondary feldspars, both albitic and anorthitic also increases. This increase in hydrothermal recrystallization is reflected by slight increase in 87Sr/S6Sr from the background ratio of 0.70268 to 0.70276 in the halo and 0.70281 for the actinolite vein. This latter ratio presumably records the Sr-isotopic composition of the reacting fluid. Actinolite however, has a low Sr content (~15ppm; Alt et al. 1996b) and neither the large increase in whole rock Sr nor the increase in 878r/86Sr can be accounted for solely by amphibole formation, which would result in a decrease in Sr content. Plagioclase is the principal reservoir of Sr within diabase (Sr ~ 1 0 0 p p m ) and the recrystallization of this phase to secondary feldspars must have a controlling influence on the Sr content and isotopic ratio of rocks from the ocean crust. Direct measurements of the Sr content or isotopic composition of secondary feldspars are

presently unavailable. Clinopyroxene contains a much lower concentration of Sr compared to plagioclase ([Sr]CP×~10ppm; Dick & Johnson 1995; Natland & Dick 1996), so the effect of clinopyroxene recrystallization on the bulk rock 87Sr/86Sr would be less than that of plagioclase alteration. Petrographic evidence indicates multiple stages of feldspar alteration, with secondary labradorite-anorthite formed early at higher temperatures and then albite-oligoclase formed at lower temperatures during later stages (Laverne et al. 1995). If the Sr isotopic compositions of hydrothermal fluids varied during the different stages, then this could affect the relative importance of these phases on the whole rock compositions. One would expect hydrothermal fluids to shift to slightly higher, more seawaterlike 87Sr/86Sr ratios with the evolution of the hydrothermal system if large fluid fluxes are involved (Bickle & Teagle 1992). Alternatively, lower, more rock-like ratios would occur if fluid fluxes are lower and fluid reacted progressively in a closed system (Alt et al. 1986a). The Sr-isotopic composition of hydrothermal minerals from the transition zone and sheeted dyke complex of Hole 504B are shown v. depth in Fig. 8. Amphiboles separated from veins have generally low 87Sr/S6Sr though they are elevated compared to the host rocks (Figs 6A & 7). Chlorite and epidote separates from the lower dykes have St-isotope ratios in the range 0.70340.7038, which is significantly higher than any whole rock sample from the lower dykes. These data indicate that the Sr-isotopic compositions of hydrothermal fluids in the lower dykes were variable and that the 87Sr/86Sr of hydrothermal fluids circulating in fractures were higher than those measured for whole rocks; a consequence of the partial recrystallization of the host diabase. High temperature (>350°C), black smokertype fluids exiting hydrothermal vents at midocean ridges have 87 Sr/ 86 Sr in the range 0.70280.7047 (Palmer & Edmond 1989), though Sr isotopic compositions can vary within a vent field (EPR 21°N, Von Damm et al. 1985) or at a single vent over time (Butterfield & Massoth 1994). The Sr contents and isotopic compositions of hydrothermal fluids are a function of reactions occurring at different temperatures along the recharge and reaction zone pathways (Berndt et al. 1988), and the range in 87 Sr/ 86 Sr ratios of different vent sites depends on the varying amounts of seawater reacting with the crust and the mineral reactions that occur, different fluid pathways, mixing of fluids in the subsurface, and differing amounts of seawater

TRACING HYDROTHERMAL FLUID EVOLUTION

91

87Sr/86Sr 0.702

0.703

0.704

0.705

0.706

600

800

E E

1000

..Q

=

iili

:

)@:

Upper Dyke Zone

.6 1200

E i

e- 1400 Q.

•

,.i-,

a Lower Dyke "Reaction"Zone

1600

\ 1800

Predicted range of 504B ~:ii,i

black-smoker fluids

Fig. 8. 87Sr/86Srof hydrothermal minerals and breccias with depth in the dyke section of Hole 504B (Kawahata et al. 1987; Alt et al. 1996a,b and Teagle unpublished data). Shaded area shows the range of the majority of whole rock samples (see Fig. 4). Stars-hydrothermal breccias; filled circles-amphibole; open circle-chlorite; filled squares-epidote; open squares-laumontite.

and hydrothermal Sr precipitated as anhydrite and epidote, respectively (Palmer & Edmond 1989). Epidote + quartz and chlorite + quartz veins cross-cut the background assemblages as well as the halos and patches in the lower dykes. Estimates from (~180 measurements indicate temperatures of formation for these minerals of up to 380°C (Alt et al. 1985a, 1986b). Abundant epidote and epidote + quartz veins are associated with sulphide mineralization in the transition zone and uppermost dykes of Hole 504B and formed at temperatures of 250-350°C. Epidotes from these veins have similar chemical compositions to, and formed over the same temperature range as the epidotes from deeper in the sheeted dykes, suggesting that they formed at a related hydrothermal stage (Alt et al. 1986a, b, 1995). The Sr-isotopic compositions of epidotes from the lower dykes are similar to epidote from a quartz + epidote vein (87Sr/86Sr=0.7038), some 650 m to 800 m shallower in the transition zone (Kawahata et al. 1987). The high temperatures of formation of hydrothermal quartz +

epidote veins in the transition zone (300-350°C; Alt et al. 1986b) limit the extent of contamination of the hot upwelling hydrothermal fluid by mixing with seawater. Epidote is commonly associated with pathways of fluid upwelling, in both in situ ocean crust as well as ophiolites (Alt et al. 1986a; Richardson et al. 1987; Vanko et al. 1992) and the 87Sr/86Sr of epidotes from the Site 504B crust most probably records the Srisotopic composition of the end-member hydrothermal fluid. Whole rock values less than 0.7034-0.7038 would then reflect variable exchange of igneous Sr in the rock with hydrothermal Sr in solution. Active black smoker vents have neither been located nor looked for at the Costa Rica Rift spreading ridge, but hydrothermal fluids with 87Sr/86Sr in the range 0.70340.7038 are consistent with the isotopic signatures of end-member hydrothermal fluids venting from a similar crustal environment along the East Pacific Rise (878r/86Sr ~0.7038+5 (lcr), Palmer & Edmond 1989). Late-stage laumontites from the transition zone have isotopic compositions within the range displayed by

92

D . A . H . TEAGLE E T A L .

A

mols SO4 per 100 m 1000

3000

B

G

87Sr/86Sr

0.702

0.704

0.706

Sr/Ca (mmol/mol) 2

0.708

4

6

8

10

12

~'~::~ Envelope of i~,,~ whole rock data

200. (w. gyrolite) "-" 400-

N

Anhydrite V e i n s

.

N:

E 600.t3 | ..Q 800-

Whole rock sulfate

E, E1000¢,.i.-, o.. iC11200-

J

Anhydrite from V e i n s and Vugs

it

.

~:

Host Diabase

~iiiiii!iiiiiiii~iiii

,q

Sulfate Leachates

1400-

¢,ol 1600-

1800-

}a

51 Range of Hole 504B "Black Smoker" fluids

Fig. 9. Distribution and chemical composition of anhydrite in Hole 504B. (A) Distribution of anhydrite in veins and pore space calculated from the occurrence of veins and the SO4 concentration of whole rock powders (Teagle et al. 1998) in 100 m depth increments per m 2 of ocean crust. (B) SVSr/86Srprofile for anhydrite with depth. Shaded zone shows the composition of carbonate and clay minerals from the uppermost volcanic rocks; C. Sr/Ca ratios for anhydrite versus depth. Shaded region shows the range of Sr/Ca ratios expected for anhydrite precipitated from mixtures of seawater and 21°N EPR black smoker fluids assuming reasonable St/Ca partition coefficients (Shikazono & Holland 1983; Berndt et al. 1988; Teagle et al. 1998). Filled squares anhydrite from vugs and veins; open squares- anhydrite leached from whole rock powders of diabase with visible anhydrite.

epidote (Kawahata e t al. 1987) though laumontite formed at lower temperatures ( < 250°C). This suggests that these hydrothermal fluids were conductively cooled during the waning stages of fluid upwelling, and not cooled by mixing with seawater.

Constraints on the evolution of recharge fluids from anhydrite A unique record of the chemical evolution of seawater during hydrothermal recharge into oceanic crust is preserved by anhydrite recovered from rocks that comprise the lower volcanic sequences and sheeted dyke complex in ODP

Hole 504B. Anhydrite (CaSO4) exhibits retrograde solubility, becoming less soluble in seaw a t e r with i n c r e a s i n g t e m p e r a t u r e , a n d precipitates from seawater at temperatures greater than > 150°C at 500 bars pressure (Bischoff & Seyfried 1978). As the molar concentration of Ca in seawater is approximately a third of that of sulphate, additional Ca must be provided from other sources in order to precipitate all the sulphate available in seawater as anhydrite. The distribution of anhydrite present in veins and as a filling of pore-space in Hole 504B is shown in Fig. 9. Anhydrite filling millimetrewide veins and precipitated in cavities is present throughout the lower volcanics, transition zone and upper sheeted dykes. The shallowest occur-

TRACING HYDROTHERMAL FLUID EVOLUTION rence of anhydrite is near the base of the upper pillow alteration zone (~300m sub-basement). Throughout the lower volcanics and transition zone, euhedral laths or radiating aggregates of anhydrite fill clay- or chlorite-lined veins. Within the upper dykes, anhydrite commonly occurs with greenschist facies minerals (chlorite, epidote), but anhydrite precipitation post-dates these minerals. At deeper levels of the dyke complex (> 1500m sub-basement) anhydrite is only rarely observed within veins and more commonly replaces igneous plagioclase in association with epidote and sodic plagioclase, or fills vugs and pore-space. Anhydrite generally does not display any obvious signs of corrosion or dissolution, but at shallow levels, anhydrite is locally partially replaced by prehnite, a zeolitelike Ca-Al-silicate (Alt et al. 1985b). The Sr-isotopic evolution of recharge fluid chemistry, as recorded by anhydrite, shows a general trend of an increasing basaltic component with depth in the crust. This matches the idealized scheme of ocean ridge hydrothermal systems with recharge via broad areas of downward fluid percolation into the crust (Fig. 9). The shallowest anhydrite encountered in Hole 504B has high 87Sr/a6Sr ratios, close to present day or 5.9 Ma seawater and very similar to Srisotopic compositions of other secondary minerals (saponite, carbonate) occurring at this and higher levels in Hole 504B and nearby Site 896 (Fig. 9). Anhydrites from the lowermost volcanic flows have high 87Sr/86Sr but with an increasing proportion of basaltic Sr. One sample from 380m sub-basement, in which bent and broken tabular prisms of anhydrite are intergrown with sheaths of gyrolite (hydrated Casilicate), yields an intermediate 87Sr/86Sr ratio (~0.7065). The unusual sulphur isotopic composition of this sample suggests a complicated hydrothermal history (see Alt et al. 1983). Basalt-fluid Sr exchange is significantly enhanced by higher temperatures and possibly lower fluid fluxes within the dykes as opposed to the more permeable pillow lavas. Anhydrites from the upper dykes have intermediate 87Sr/S6Sr (~0.7045-0.7056) and there is a general trend of an increasing basaltic component with depth. Very little vein anhydrite is recovered below 1200m sub-basement and the recharge profile for the lowermost sheeted dykes was completed by analysing anhydrite leached from powders of whole rocks that contain significant sulphate (>200 ppm SO4), or have visible anhydrite within pore spaces (see Teagle et al. 1998). These samples yield low, basalt-like 87Sr/86Sr. Two samples have strontium isotopic compositions nearly identical to the host basalts,

93

whereas the other two samples have significant, albeit minor, seawater derived Sr. The anhydrite measurements from the lower sheeted dykes are within the range of 87Sr/S6Sr compositions of secondary greenschist minerals in the rocks (actinolite, epidote, chlorite; 87Sr/86Sr ~0.7028-0.7038), but the lowest values are less than the predicted composition of upwelling black smoker-type fluids as recorded by epidote (87Sr/S6Sr~0.7034-0.7038). A single analysis of anhydrite from the mineralized stockwork has a Sr-isotopic composition identical to epidote (STSr/a6Sr~0.7038) from the same sample. Shikazono & Holland (1983) have measured the Sr/Ca partition coefficient for anhydrite to be in the range 0.27 to 0.73 following KSr: d

(asrsO4~/(asr2+~ \ acaso4 , / / k , aca2+ ,/

(1)

and the Sr/Ca ratio of anhydrite can be used to infer the Sr/Ca ratio of fluid from which it precipitated (Teagle et al. 1998). The downward decrease of anhydrite Sr/Ca ratios in the volcanic sequences (Fig. 9) indicates decreasing Sr concentrations in the fluids resulting from the precipitation of calcium carbonate and anhydrite. These fluids retain near-seawater STSr/86Sr indicating only minor involvement of basaltic Sr, which probably precludes a significant increase in fluid Ca concentration (Seyfried & Mottl 1982). Greater fluid-rock reaction occurred in the lowermost pillow lavas and upper sheeted dykes as recorded by the formation of albite-oligoclase and chlorite, and a concomitant shift to intermediate STSr/86Sr and very low Sr/ Ca ratios in anhydrite from these levels. Experimental studies indicate that Sr is more strongly partitioned into albite than Ca (Blundy & Wood 1991) and albitiztion of the upper dykes may account for the very low Sr concentration of the recharge fluids as suggested by anhydrite. Strontium exchange between the recharge fluids and rocks increases with depth in the sheeted dyke complex, though the low Sr/Ca ratios from these deep levels suggest that fluid Sr concentrations remain uniformly low. The final step in the evolution of seawater to black smoker-type fluids probably involves the formation of secondary calcic plagioclase (Berndt et al. 1989; Berndt & Seyfried 1993) which is found in the lower dykes of Hole 504B, but this process is not constrained by the anhydrite data.

Summary The Sr-isotopic signature of basalts and diabases

94

D.A. H. TEAGLE E T AL.

from the upper oceanic crust, is dependent on the proportions of secondary minerals and in particular the extent of recrystallization of igneous plagioclase, the principal St-bearing phase. Igneous plagioclase has remained unaltered or only partially altered in the background rocks throughout the upper crust, resulting in whole rock 87Sr/86Sr ratios intermediate between fluid and igneous compositions, and indicating fluid-rock Sr-isotopic disequilibrium. Highly altered rocks, such as hyaloclastic and hydrothermal breccias and sub-greenschist facies basalts from the transition zone and upper dykes, comprise nearly 100% secondary minerals and these rocks preserve basement fluid compositions. Sr-isotopic analyses of whole rock samples record the decreasing influence of seawater Sr with depth in the ocean crust. Secondary minerals also record the composition of fluids percolating into the crust. Fluid circulation in the uppermost crust ( < 300 m subbasement), was initiated at the ridge axis and continues today. Smectite and carbonates formed from fluids with near-seawater Srisotopic compositions but containing a minor basaltic component (87Sr/86Sr ~0.708-0.709; 520% MORB-derived Sr). The large fluid fluxes associated with low temperature circulation on the ridge flanks means that even minor changes in fluid chemistry due to interaction with the uppermost basement could have significant effects on the oceanic chemical budgets, in particular the global strontium budget (Palmer & Edmond 1989). Anhydrite, present throughout the lower pillow lavas, transition zone and dyke complex, records the evolution of basement fluids from seawater-like compositions to Sr-isotopic signatures identical to hydrothermal minerals preserved at the bottom of the dyke section. Only limited fluid modification occurs down to the bottom of the pillow lavas, with anhydrite compositions similar to the clay minerals and carbonates in the overlying pillow lavas. The change to intermediate and then MORB-like Srisotopic signatures occurs rapidly with depth in the dykes (87Sr/86Sr 0.7055->0.7030), and reflects the higher temperature of reaction and the greater extent of plagioclase recrystallization to more sodic compositions in the upper dykes. The isotopic composition of basement fluids deep in the dyke complex is preserved by amphibole, which is the principal secondary mineral developed at these levels. Amphibole analyses indicate that there was a range of hydrothermal fluid compositions responsible for the alteration of the lower dykes. Vein minerals have higher SVSr/S6Sr compared to their asso-

ciated alteration halos and background host diabase, reflecting the partial replacement of primary igneous phases, but fluid compositions were strongly buffered towards rock dominated 87Sr/86Sr (~0.7027~0.7030). Although anhydrite is generally late in the paragenetic sequence preserved in the lower dykes, the lowest 87Sr/86Sr anhydrite analyses from this section of the crust are within the range of compositions recorded by amphibole, and may record a recharge fluid evolution appropriate for the background greenschist facies alteration observed in the lower sheeted dykes. Analyses of epidote and chlorite from hydrothermal veins that cross-cut the amphibole veins, associated alteration halos and the background dark grey diabase yield a limited range of S7Sr/S6Sr (~0.7034-0.7038). Similar Sr-isotopic compositions are preserved by epidote and laumontite from the transition zone 600 to 800m shallower in the crust and suggest that upwelling black smoker-type fluids in the 504B crust had SYSr/86Sr ~0.7034-0.7038. This range is higher than the 87Sr/86Sr of the most evolved deep fluids recorded by both amphibole and anhydrite. The very low Sr/Ca ratios of anhydrite from the dykes requires fluids more depleted in Sr than any black smoker fluid recovered to date (for reasonable estimates of fluid-anhydrite partition coefficient). Hence, the black smoker fluids that precipitated epidote, quartz and chlorite in veins in the lower dykes and transition zone of Hole 504B, must have reacted elsewhere in the crust, and less Sr exchange occurred during recharge for these fluids. The higher 87 Sr/ 86 Sr ratios of epidote, and by inference the Hole 504B black smoker fluids, must reflect localised higher fluid fluxes resulting from the channelling of recharge fluids. The 'reaction zone' where fluids evolve to black smoker-type compositions must migrate within the lower sheeted dyke complex and uppermost plutonic rocks depending on the interplay between magmatic and hydrothermal processes, in particular the local permeability structure. Such reactions are probably recorded by the formation of secondary calcic plagioclase rims in the more highly altered patches and vein halos, which record an early high temperature phase of alteration in the lower sheeted dykes of Hole 504B. It is predicted that the secondary plagioclases from the lower dykes should yield 87Sr/86Sr similar to the range displayed by epidote. As hydrothermal recharge into the lower dykes becomes more restricted due to the clogging of fluid pathways by secondary minerals, basement fluids will evolve from 'black smoker'-like to more rock dominated Sr-isotopic

TRACING HYDROTHERMAL FLUID EVOLUTION signatures as recorded by amphiboles ( ~ 0 . 7 0 2 8 0.703). This research was supported by grants to J. C. Alt and D. A. H. Teagle from the National Science Foundation (OCE-9217423, OCE-9314218 & 9633773), and JOI/USSAC (USSSP 148-20731b & USSSP 14820738b). The authors would like to thank T. Huston and J. N. Christensen for analytical assistance. The manuscript greatly benefited from discussion with D. Vanko and the two anonymous reviewers are thanked for their constructive criticism and advice. D.T.'s attendance of the Geological Society of London meeting (May 1997) was supported by the Sokol Fellowship of the University of Michigan. References

ALa-, J. C. 1995. Subseafloor processes in Mid-Ocean Ridge Hydrothermal Deposits. In." HuMwmS, S. E., ZIERENBERG, R. A., MULLINEAUX, L. S. & THOMSON, R. E. (eds) Physical, Chem&al, Biological and Geological Interactions within Submarine Hydrothermal Systems. Geophysical Monograph, 91, Washington, AGU, 85-114. - & HONNOREZ,J. 1984. Alteration of the upper oceanic crust DSDP Site 417: mineralogy and chemistry. Contributions to Mineralogy and Petrology, 87, 149-169. --, HtmBER'rEN, H.-W. & SALrZMAN, E. 1983. Occurrence and origin of anhydrite from Deep Sea Drilling Project Leg 70, Hole 504B, Costa Rica Rift. In." CAr~, J. R., LAN~SFH, M. G. & HOYNOREZ, J. (eds) Initial Reports of the Deep Sea Drilling Project, 69, Washington, U.S. Govt Printing Office, 547 550. , LAVERNE, C. & MUEHLENBACHS, K. 1985a. Alteration of the upper oceanic crust: mineralogy and processes in Deep Sea Drilling Project Hole 504B, Leg 83. In: ANDERSON,R. N., HONNOREZ,J. & BECWER,K. (eds) Initial Reports of the Deep Sea Drilling Project. 83, Washington, U.S. Govt Printing Office, 217-247. , SALTZMAN, E. S. & PRICE, D. A. 1985b. Anhydrite in hydrothermally altered basalts: Deep Sea Drilling Project Hole 504B, Leg 83. In: ANDERSON,R. M., HONNOREZ,J. & BECKER,K. (eds) Initial Reports of the Deep Sea Drilling Project, 83, Washington, U.S. Govt Printing Office, 283-288. , HONNOREZ,J., LAVERNE,C. & EMMERMANN,R. 1986a. Hydrothermal alteration of a I km section through the upper oceanic crust, Deep Sea Drilling Project Hole 504B: mineralogy, chemistry, and evolution of seawatewbasalt interactions. Journal of Geophysical Research, 91, 1030910335. , MUEHLENBACHS,K. & HONNOREZ,J. 1986b. An oxygen isotope profile through the upper kilometre of the oceanic crust, DSDP Hole 504B. Earth and Planetary Science Letters, 80, 217-229. --, ANDERSON, T. F. & BONNELL, L. 1989. The geochemistry of sulfur in a 1.3km section of

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hydrothermally altered oceanic crust, DSDP Hole 504B. Geochimica et Cosmochimica Acta, 53, 1011-1023. , KINOSHITA,H., STOKKINa, L. B., et al. 1993. Proceedings of the Ocean Drilling Program, Initial Reports, 148, College Station, TX (Ocean Drilling Program), 352pp. , Zuleger, E. & Erzinger, J. A. 1995. Mineralogy and stable isotopic compositions of the hydrothermally altered lower sheeted dike complex, Hole 504B, Leg 140. In. ERZINGERJ., BECKER,K., D1cg:, H. J. B. & S'rOKKING,L. B. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 137/140, College Station, TX (Ocean Drilling Program), 155-166. - - , LAVERNE, C., VANKO, D. et al. 1996a. Hydrothermal alteration of a section of upper oceanic crust in the Eastern Equatorial Pacific: A synthesis of results from DSDP/ODP Legs 69, 70, 83, 111,137, 140, and 148 at Site 504B. In: ALT, J. C., KINOSHITA,H., STOKKING,L. B. & MICHAEL,P. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, College Station, TX (Ocean Drilling Program), 417-434. , TEAGLE,D. A. H., BACH, W., HALLIDAY,A. N. & ERZIN~ER, J. 1996b. Stable and strontium isotopic profiles through hydrothermally altered upper oceanic crust, ODP Hole 504B. In: ALT J. C., KINOSHITA,H., STOKKING,L. B. & MICHAEL,P. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, College Station, TX (Ocean Drilling Program), 57-70. , LAVERNE, C., VANKO, D., BACH, W., HONNOREZ, J. & BECr~ER, K. 1996c. Ridge flank alteration of upper oceanic crust in the Eastern Pacific: a synthesis of results for volcanic rocks of Holes 504B and 896A. In." ALT, J. C., KINOSHITA, H., STOKKING, L. B. & MICHAEL, P. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, College Station, TX (Ocean Drilling Program), 435-452. BARRETr, T. J. 1983. Strontium- and lead-isotope composition of some basalts from Deep Sea Drilling Project Hole 504B, Costa Rica Rift, Legs 69 and 70. In: CANN, J. R., LANaSE'rH, M. G., HONNOREZ,J. (eds) lnitial Reports of the Deep Sea Drilling Project, 69, Washington, U.S. Govt. Printing Office, 643-650. & FRIEDRICHSEN, H. 1982. Strontium and oxygen isotopic composition of some basalts from Hole 504B, Costa Rica Rift, DSDP Legs 69 and 70. Earth and Planetary Science Letters, 60, 27-38. - &- 1987. Oxygen-isotopic compositions of basalts from young spreading axes in the eastern Pacific. Canadian Journal of Earth Sciences, 24, 2105-2117. BECKER, K. & SA~I, H., et al. 1988. Proceeding of the Ocean Drilling Program, Initial Reports, 111, College Station, TX (Ocean Drilling Program), 357pp. , ADAMSON, A. ET .4L. 1989. Drilling deep into young oceanic crust, Hole 504B, Costa Rica Rift. Reviews in Geophysics, 27, 79-102. BERND% M. E. & SEYERIED,W. E. 1993. Calcium and

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sodium exchange during hydrothermal alteration of calcic plagioclase at 400°C and 400 bars. Geochimica et Cosmochimica Acta, 57, 4445-4451. , SEYFR1ED, W. E. & BECK, J. W. 1988. Hydrothermal alteration processes at mid-ocean ridges: experimental and theoretical constraints from Ca and Sr exchange reactions and Sr isotope ratios. Journal of Geophysical Research, 93, 45734583. , & JANECKY, D. R. 1989. Plagioclase and epidote buffering of cation ratios in midocean ridge hydrothermal fluids: Experimental results in and near the super-critical region. Geochirnica et Cosmochimica Acta, 53, 2283-2300. BICKLE, M. J. & TEAGLE, D. A. H. 1992. Strontium alteration in the Troodos ophiolite: implications for fluid fluxes and geochemical transport in midocean ridge hydrothermal systems. Earth and Planetary Science Letters, 113, 219-237. BISCHOFE, J. L. & SEYFR~ED,W. E. 1978. Hydrothermal chemistry of seawater from 25°C to 350°C. American Journal of Science, 278, 838-860. BLUNDY, J. D. & WOOD, B. J. 1991. Crystal-chemical controls on the partitioning of Sr and Ba between plagioclase feldspar, silicate melts, and hydrothermal solutions, Geochimica et Cosmochimica Acta, 55, 193-209. BREWER, T. S., HARVEY, P. K., LOVELL, M. A. & WILLIAMSON, G. 1995. Stratigraphy of the ocean crust in ODP Hole 896A from FMS inqages. Scientific Drilling, 5, 87-92. BURKE, W. H., DENISON, R. E., HETHERINGTON,E. A., KOEPNICK, R. B., NELSON, H. F. & OTIO, J. B. 1982. Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology, 10, 516-519. BUTTERFIELD, D. A. & MASSOTH, G. J. 1994. Geochemistry of north Cleft segment vent fluids: Temporal changes in chlorinity and their possible relation to recent volcanism. Journal of Geophysical Research, 99, 4951-4969. DICK, H. J. B. & JOHNSON, K. T. M. 1995. REE and trace element composition of clinopyroxene megacrysts, xenocrysts, and phenocrysts in two diabase dikes from Leg 140, Hole 504B. In: ERZINGER, J., BECKER, K., DICK, H. J. B. & STOKKING, L. B. (eds) Proceeding of the Ocean Drilling Program, Scientific Results, 137/140, College Station, TX (Ocean Drilling Program), 121-130. FISHER, A. T., BECKER,K. & NARASIMHAN,T. N. 1994. Off-axis hydrothermal circulation: Parametric tests of a refined model of processes at DSDP/ ODP Site 504. Journal of Geophysical Research, 99, 3097-3123. , - - , LANGSETH, M. G. & MOTTL, M. J'. 1990. Passive off-axis convection through the southern flank of the Costa Rica Rift. Journal of Geophysical Research, 95, 9343-9370. FRIEDRICHSEN, H. 1985. Strontium, oxygen, and hydrogen isotope studies on primary and secondary minerals in basalts from the Costa Rica Rift, Deep Sea Drilling Project Hole 504B, Leg 83. In." ANDERSON,R. N., HONNOREZ,J., BECKER,K. (eds) Initial Reports of the Deep Sea Drilling Project 83,

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Washington, U.S. Govt Printing Office, 289-295. HART, S. R. & MOTTL, M. J. 1983. Alkali and Sr isotope geochemistry of waters collected from basaltic basement, Deep Sea Drilling Project Hole 504B, Costa Rica Rift. In." CANN, J. R., LANGSETH, M. G., HONNOREZ, J. (eds) Initial Reports of the Deep Sea Drilling Proiect 69, Washington, U.S. Govt Printing Office, 487-494. --, BLUSZTAJN,J., DICK, H. J. B. & LAWRENCE,J. R. 1994. Fluid circulation in the oceanic crust: contrast between volcanic and plutonic regimes. Journal of Geophysical Research, 99, 3163-3174. HODELL, D. A., MUELLER, P. A. & GARRIOO, J. R. 1991. Variation in the strontium isotopic composition of seawater during the Neogene. Geology, 19, 24-27. KAWAHATA, H., KUSAKABE,M. & KIKUCHI, Y. 1987. Strontium, oxygen and hydrogen isotope geochemistry of hydrothermally altered and weathered rocks in DSDP Hole 504B, Costa Rica Rift. Earth and Planetary Science Letters, 85, 343-355. KUSAKABE, M., SHIBATA, T., YAMAMOTO, M. E T A L . 1989. Petrology and isotopic characteristics (H, O, S, Sr, and Nd) of basalts from Ocean Drilling Program Hole 504B, Leg 111, Costa Rica Rift. In: BECKER, K. & SAKAI, H. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 111, College Station, TX (Ocean Drilling Program), 47 60. LANGSETH, M. G., MOTTL, M. J., HOBART, M. & FISHER, A. 1988. The distribution of geothermal and geochemical gradients near Site 501/504. In." BECKER K. & SAKAI, H. (eds) Proceedings of the Ocean Drilling Program, Initial Reports, 111, College Station, TX (Ocean Drilling Program), 23 32. LAVERNE,C., VANKO,D. A., TARTAROTTI,P. & ALT, J. C. 1995. Chemistry and geothermometry of secondary minerals from the deep sheeted dike complex, Hole 504B. In: ERZINGER, J., BECKER, K., DICK, H. J. B. & STOkKING, L. B. (eds) Proceedings of the Ocean Drilling Program, Scient(fic Results, 137/140, College Station, TX (Ocean Drilling Program), 167-189. LILLEY, M. D., BUTTERFIELD, D. A., OLSON, E. J., LUPTON, J. E., MACKO, S. A. & MCDUFF, R. E. 1993. Anomalous CH4 and NH4 ÷ concentrations at an unsedimented mid-ocean ridge hydrothermal system. Nature, 364, 45-47. MOTTL, M. J. 1989. Hydrothermal convection, reaction and diffusion in sediments on the Costa Rica Rift Flank: pore-water evidence from ODP Sites 677 and 678. hT.' BECKER, K. & SAKAI, H. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 111, College Station, TX (Ocean Drilling Program), 195-213. & WHEAT, C. G. 1994. Hydrothermal circulation through mid-ocean ridge flanks: fluxes of heat and magnesium. Geochimica et Cosmochimica Acta, 58, 2225-2238. NATLAND,J. H. & DICK, H. J. B. 1996. Melt migration through high-level gabbroic cumulates of the East Pacific Rise at Hess Deep: The origin of magma lenses and the deep crustal structure of fast-

TRACING HYDROTHERMAL FLUID EVOLUTION spreading ridges. In. M~VEL, C., GILL/S, K. M., ALLAN, J. F. & MEYER, P. S. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 147, College Station, TX (Ocean Drilling Program), 21-58. PALMER,M. R. & EDMOND,J. M. 1989. The strontium isotope budget of the modern ocean. Earth and Planetary Science Letters, 92, 11-26. RICHARDSON, C. J., CANN, J. R., RICHARDS,H. G. & COWAN, J. G. 1987. Metal-depleted root zones of the Troodos ore-forming hydrothermal systems, Cyprus. Earth and Planetary Science Letters, 84, 243 253. SEYFRIED,W. E. & MOTTL, M. J. 1982. Hydrothermal alteration of basalt by seawater under seawaterdominated conditions, Geochimica et Cosmochimica Acta, 46, 985-1002. SHIKAZONO, N. & HOLLAND, H. D. 1983. The partitioning of strontium between anhydrite and aqueous solutions from 150 ° to 250°C. In." OHMOTO, H. & SKINNER, B. J. (eds) The Kuroko and Related Volcanogenic Massive Sulfide Deposits. Economic Geology Monograph, 5, 320-328. SHIMIZU, H., MORI, K. & MASUDA,A. 1989. REE, Ba and Sr abundances and Sr, Nd, and Ce isotopic ratios in Hole 504B basalts, ODP Leg 111, Costa Rica Rift. In: BECKER, K. & SAKAI, H. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, I l l , College Station, TX (Ocean Drilling Program), 77-83. SPOONER, E. Z. C., CHAPMAN,H. J. ~g SMEWING~J. D. 1977. Strontium isotopic contamination and oxidation during ocean floor hydrothermal metamorphism of ophiolitic rocks of the Troodos Massif, Cyprus. Geochimica et Cosmochimica Acta, 41, 873-890. STAUDIGEL, H. & HART, S. R. 1985. Dating of ocean crust hydrothermal alteration: Strontium isotope ratios from Hole 504B carbonates and the reinterpretation of Sr isotope data from Deep Sea Drilling Project Sites 105, 332, 417, and 418. In: ANDERSON,R. N., HONNOREZ,J., BECKER,K. (eds) Initial Reports of the Deep Sea Drilling Project, 83, Washington, U.S. Govt Printing Office, 297-303.

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, DAVIS, G. R., HART, S. R., MARCHANT,K. M. SMITH,B. M. 1995. Large scale isotopic Sr, Nd and O isotopic anatomy of altered oceanic crust: DSDP/ODP sites 417/418. Earth and Planetary Science Letters, 130, 169-185. STEIN, C. A. & STEIN, S. 1994. Constraints on hydrothermal heat flux through the oceanic lithosphere from global heat flow. Journal of Geophysical Research, 99, 3081-3096. TEAGLE, D. A. H., ALT, J. C., BACH,W., HALLIDAY,A. N. & ERZINGER, J. 1996. Alteration of upper oceanic crust in a ridge flank hydrothermal upflow zone: mineral, chemical and isotopic constraints from ODP Hole 896A. In: ALT, J. C., KINOSHITA,H., STOKKING,L. B. & MICHAEL,P. J. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, College Station, TX (Ocean Drilling Program), 119-150. & HALLIDAY,A. N. 1998. Tracing the chemical evolution of fluids during hydrothermal recharge: Constraints from anhydrite recovered in ODP Hole 504B. Earth and Planetary Science Letters, 155, 167-182. VANKO, D. A., GRIFFITH,J. D. & ERICKSON,C. L. 1992. Calcium-rich brines and other hydrothermal fluids in fluid inclusions from plutonic rocks, Oceanographer Transform, Mid-Atlantic Ridge. Geochimica et Cosmochimica Acta, 56, 35-47. --, LAVERENE,C., TARTAROTTI, P. & ALT, J. C. 1996. Chemistry and origin of secondary minerals from the deep sheeted dikes cored during Leg 148, Hole 504B. In: ALT, J. C., KINOSHITA, H., STOKKING, L. B. & MICHAEL, P. J. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, College Station, TX (Ocean Drilling Program), 71-86. VON DAMM, K. L., EDMOND, J. M., GRANT, B., MEASURES, C. I., WALDEN, B. & WEISS, R. F. 1985. Chemistry of hydrothermal solutions at 21°N, East Pacific Rise. Geochimica et Cosmochimica Acta, 49, 1031-1083. WELHAN, J. A. t~ CRAIG, H. 1979. Methane and hydrogen in East Pacific Rise hydrothermal fluids. Geophysical Research Letters, 6, 829-831. -

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Petrological investigations of low temperature hydrothermal alteration of the upper crust, Juan de Fuca Ridge, ODP Leg 168 A. G. H U N T E R

1 & O D P L E G 168 S C I E N T I F I C

PARTY 2

1 School o f Earth Sciences, University o f Leeds, Leeds L S 2 9JT, U K 2 c/o Ocean Drilling Program, Texas A & M University, 1000 Discovery Drive, College Station, T e x a s 77845-9547, U S A Abstract: Ten sites were drilled across the eastern flank of the Juan de Fuca Ridge during

ODP Leg 168 to investigate the nature of ridge~ank hydrothermal circulation and crustal evolution in a region of varying sediment-basement topography. The sites were divided into: (i) the Hydrothermal Transition (HT) transect, a zone of gradually thickening sedimentary cover, extending from the ridge-flank boundary for a distance of 20 km to the east; (ii) the Buried Basement (BB) transect, which consists of uniformly, flat lying basement, capped by a regionally continuous thick sedimentary sequence; and (iii) the Rough Basement (RB) transect, represented by a region of basement highs and troughs, located ~100km east of the ridge axis, and capped by a variable thickness of sediment. Across the flank, a progressive sequence of alteration was recognized, produced as the hydrothermal system evolved from open oxidative conditions involving a slightly modified seawater derived fluid, to a more restricted/closed, non-oxidative environment in which more strongly modified seawater derived fluids interacted with the basement. Alteration commenced with the formation of chlorite at relatively high temperatures in non-oxidative to oxidative conditions, followed by the first wide spread stage of low temperature oxidative alteration with the formation of iron oxyhydroxides and celadonite. Progressive burial, along with continued alteration of the basement, restricted fluid circulation resulting in the formation of saponite dzpyrite over a range of temperatures under non-oxidative conditions. The final stage of alteration is represented by carbonate, which formed under non-oxidative conditions by interaction of strongly modified seawater and the crust.

Mid-oceanic ridges are i m p o r t a n t sites of advective heat loss and water-rock interaction, and as such have formed the basis of many studies over the last few decades. Simple modelling along with direct observations (e.g. Fisher et al. 1990; Langseth et al. 1992; Davis & Currie 1993) however, have shown that more than three times the amount of heat is lost and more than ten times the amount of fluid fluxes through ridge flanks compared to ridge axes (Davis et al. 1992). Advective heat loss by interaction between hydrothermal fluids and the oceanic crust is known to continue for several tens of millions of years (Anderson et al. 1977). This implies that at present, more than one third of the oceanic crust is actively involved in some form of h y d r o t h e r m a l interaction (Anderson et al. 1977), altering the primary mineralogy, geochemistry and physical properties of the upper crust. However, due to the wide range of conditions found within and across ridge flanks (e.g. rock type, permeability structures, depth of burial), as well as in hydrot h e r m a l fluids (e.g. c o m p o s i t i o n , salinity, temperature, fluid flow rate), the range of

controls affecting water-rock interaction in this environment have yet to be fully quantified. To date, previous studies on this topic (e.g. Bass 1976; Andrews 1977; B6hlke et al. 1980; Honnorez et al. 1983; A l t e t al. 1986, 1992, 1996; Adamson & Richards 1990; Laverne et al. 1996; Teagle et al. 1996) have highlighted several important controls, including: the water-rock temperature; the rate of fluid flux; the primary texture and mineralogy of the basement; the method of fluid-basement interaction; and the presence/absence of sediment, all of which will influence how the oceanic crust is affected by low temperature hydrothermal alteration. In an attempt to quantify and narrow the range of potential factors affecting low temperature hydrothermal alteration of the ocean crust, ODP Leg 168 drilled a series of holes across the eastern flank of the Juan de Fuca Ridge, where a number of different hydrothermal regimes had been recognized during previous submersible and geophysical surveys (Davis et al. 1992; Wheat & Mottl 1994; Thompson et al. 1995). By carrying out detailed hydrological, geochemical, petrological and geophysical stu-

HUNTER. A. G. & ODP LEG 168 SCIENTIFICPARTY. 1998. Petrological investigations of low temperature hydrothermal alteration of the upper crust, Juan de Fuca Ridge, ODP Leg 168 In: MILLS,R. A. • HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 99-125

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A. G. HUNTER E T AL.

dies on the oceanic crust, sedimentary units and pore fluids, the primary objective of ODP Leg 168 was to enable a more thorough understanding of the magnitude of fluid--crust chemical exchange, and to establish what the primary factors controlling water-rock interaction, changes in fluid chemistry and the physiochemical alteration paths are in the oceanic crust. Petrological and microprobe analyses on secondary mineral phases from samples from ODP Leg 168 are presented in this paper, with the primary objectives of this study to: (i) constrain the extent, type and sequence of alteration in the upper oceanic crust; and (ii) to establish how this has been affected by variations in lithology, hydrothermal fluid composition and basement-sediment relationships.

ODP Leg 168: eastern flank of the Juan de Fuca Ridge-geological setting The Juan de Fuca Ridge (JdFR) is situated ~400 km off the Northwest American coast (Fig. 1), and is spreading at a half rate of c. 58 mma -1 (Davis & Currie 1993). Beneath the flat lying

turbiditic sedimentary cover, the basement topography is dominated by a series of linear ridges and troughs, formed by normal block faulting and variations in the magma supply rate (Davis et al., 1997). Previous sea-surface and shallow drilling expeditions (Davis et al. 1992; Wheat & Mottl 1994; Thompson et al. 1995) have identified several distinct hydrothermal regimes across the eastern flank, controlled by a combination of different physical factors including basement relief and sediment cover (Fig. 2a). During ODP Leg 168, ten sites (1023-1032) were drilled on the eastern flank of the Endeavour segment in a linear transect perpendicular to the Juan de Fuca Ridge (Figs l b & 2a). The sites were divided into three transects characterized by different hydrothermal regimes, and are referred to as: (i) the Hydrothermal Transition (HT) transect, located 20 km to the east of the ridge axis; (ii) the Buried Basement (BB) transect, commencing c. 40 km east of the ridge axis; and (iii) the Rough Basement (RB) transect, commencing c. 90km east of the ridge axis (Davis et al., 1997; Fig. 2a, b).

Plate

d

J u a n de Fuca Plate

Fig. 1. (a) Location map of ODP Leg 168 (represented by the striped line) and the eastern flank of the Juan de Fuca Ridge, NE Pacific Ocean;

PETROLOGY OF LOW TEMPERATURE ALTERATION

101

48 ff5'N

48 ~O0'N

47°45'N

47~30'N 129 °W

128 °W

Q Hydrothermal Transition Transect (HT)

4.0

• Buried Basement • Rough Basement ,Transect (BB) ......... Transect (RB) .....

©

Site 1 0 2 7 , 3.59~

....

3.5

//Site 1026, 3.49Ma

/

3,0

Site t 0 3 2 , " / / ~ 2.62Ma~l t ~,~1 2.0 ~

.1 Site 1029, Site t028, f 1.95Ma 1.6oM,~i Site 1025, /1Sites 1030& 1031, 1.27ya~ 1.46Ma Site 1 0 2 3 ~ Site 1024, 0 ~ 0.99Ma

1.5 1.0 0.5

/ 0.0

I

0

20

I

I

I

I

40

60

80

100

120

Distance f r o m the ridge axis (km)

Fig. 1. (b) Magnetic anomaly map of the ten sites drilled during ODP Leg 168, and the division of the three hydrothermal transects-the Hydrothermal transition (HT) transect, the Buried Basement (BB) transect and the Rough Basement (RB) transect (Adapted from Davis et al. 1997); (e) Correlation between distance from the ridge axis and inferred crustal age for Sites 1023-1032. Inferred crustal ages are from Cande & Kent, 1992; distances are according to Davis (pers. comm.) (Adapted from Davis et al. 1997.)

The following section briefly describes the diagnostic features of the basement, overlying sediments and hydrothermal regime in the three transects; a detailed account can be obtained from the ODP Initial Reports volume 168 (Davis et al., 1997).

Hydrothermal transition ( H T ) sites 1023-1025

transect,

Approximately 20 km east of the ridge axis of the Endeavour segment of the JdFR, there is a sudden change in basement relief, marking the

. LL~ r~

~km

--

30

~tteration Phases !, chlorite-smectite i, ± celadonite+ iddingsite : saponite+ pyrite ± minorcarbonate, zeolite, quartz ,.-:-.,~:~"

i 20

40 °

,

t

1.5 51 °

10

1025 1030-31 1028

I I I

I

I I I

380

i

50

60

70

Eastern Flank

57 °

,

I

. . . .

1032

Alteration Phases • celadonite+ iddingsite saponite + pyrite • carbonate+ minorzeolite,talc •

t029

.59 °

Buried Basement (BB) 2.0 2.5 , I I

80

3.0 I

I lOOkm

>

iAlteration Phases i• chlori~e-smectite ~-eeladonite+ iddingsite i" safronite+ pyrite ~. carbonate+ i'.eolite

90

1026 1027

I

!

61-64 °

3.5Ma

Rough Basement (RB)

Fig. 2. (a) Schematic cross-section of the eastern flank of the Juan de Fuca Ridge (derived from seismic reflection) illustrating the changing topography of the sedimentary units and underlying igneous basement. The relative locations of Sites 1023-1032 are indicated, with measured basement-sediment temperatures and inferred basement ages (Davis et al. 1997). A summary of the main alteration mineral assemblage present in the basement is listed for all three transects. The black arrows represent the proposed major fluid flow directions in the flank, based on direct fluid flow measurements, changes in the fluid chemistry and changes in the physical properties of the =rust, obtained during ODP Leg 168 (adapted from Davis et al. 1997).

I0

1024

Site! 1023

~

Hydrothermal Transition (liT) !.0 22 o

0.5

i

"Basement-Sediment Interft'we ;6 ° Temperature (°C)

0.0 O.

~ - ................Ridge Axis

Age

Ridge Axis

PETROLOGY OF LOW TEMPERATURE ALTERATION Hydrothermal Transition 1023

1024

: :i::,~:,

~i~::::::::

102,5

.:i:: ...

100

1031

Buried Basement t028 1029

103 Rough Basement

t032

1026

1027

o ,,] 1oo ~

~ 200

"

R

k

t t

@

2oo~

R 30O

L

iiiii;

!

300

R L

~4oo

Recovery Log* Lithology Log

!!;~:i!

R

....

L

400

400

5ooJ

500

600,

6OO

700.

700

Sediment

Breccia

500 ¸

B

N

Pillowb~atts Massive basaIts

600,

l 700

Doteritesheet

R

*Recoveryshownfor igneous basementonly; 100% recoveryassumedfor sediments "

L

Fig. 2. (b) Composite lithostratigraphical logs for the HT, BB and RB transects, graphically illustrating the varying thicknesses of basement cored at each site. Column R represents the combined recovery for all holes in each site across the three transects, indicated graphically by the black horizontal lines of variable thickness; column L represents a composite of the lithological record for all holes in each site. The lithological units are shown so that recovery (always < 100%), has been expanded to fit the interval cored. boundary between the elevated ridge axis and western margin of the Cascadian Basin, i.e. the ridge flank (Fig. 2a). Associated with this, is a sharp change from virtually sediment free basement at the ridge axis, to basement which is blanketed by a regionally continuous thick sequence of hemipelagic to pelagic turbidites. This change in sediment cover has a fundamental effect on hydrothermal circulation, with the ridge axis undergoing open circulation and free exchange between the seawater and crust, whilst sedimentary cover on the ridge flanks restricts exchange between the circulating oxidative seawater-derived fluids and crust, throughout the ridge flank. Changes in the permeability structure of the crust between the ridge axis and flank will also play an important role in influencing hydrothermal fluid flow. In addition to changes in fluid flow, the boundary between the ridge axis and flank is also marked by sharp rises in basement-sediment interface temperature and conserved lithospheric heat flow, and by distinct changes in fluid composition (Davis et al. 1997). These physiochemical changes have been correlated with a shift in hydrothermal circulation from open (i.e. ridge axis, with free exchange between the fluid and crust) to restricted (i.e. ridge flank, with limited flui& crust chemical exchange), and is referred to as the Hydrothermal Transition (HT) transect

(Fig. 2a). The HT sites (1023-1025), intersect basement varying in age from 0.86-1.27 Ma (Fig. lb~:), with the sediment-basement interface temperature increasing from c. 10°C near the ridge axis to 15°C at the axis-flank boundary and 38°C at ~20 km east of this boundary (Davis et al. 1992; Davis et al. 1997; Fig. 2a). Over the same distance, the underlying lithospheric heat flow increases from 15% to > 8 0 % of the expected conductive value (Davis et al. 1992; Davis et al. 1997). Across the transect, igneous crustal basement penetration (below the sedimentary units) varied between sites from < 5m (Sites 1023 and 1024) to ~41 m (Site 1025; Fig. 2b; Davies et al. 1997). Buried Basement (BB) 1032

transect, sites 1 0 2 8 -

The Buried Basement (BB) Transect, spans a region c. 40-100 km east of the ridge axis (Figs 1b, c & 2a), and is characterized by a relatively smooth basement topography, covered by regionally continuous hemipelagic to pelagic turbiditic muds and clays (Davis et al. 1992; Davis et al. 1997). Within this transect, local advective heat loss and water-rock exchange are restricted by the thick and uniform sedimentary cover. The transect covers crust that varies in

1% zeol 1%

Trace

Trace

< 1%

-

Trace

Primary phenocrysts Plag. O1.

-

-

-

Px.

sap-py-zeol

sap-chl-py-carb-

sap-cel-py

cel-idd-sap-py

5

5 5 2 1 3

Secondary alteration Phases %

Pervasive

Rim Core Rim Core Pervasive

Analysis

0.702534

0.702470 0.702478 0.702486 0.702456 0.702787

87Sr/86Srn

6.79

7.37 6.69 6.80 6.65 703

(5t80

0.25

0.13 0.15 0.16 0.11 0.61

W/Rsr (open)

Secondary phases are: carb: carbonate; cel: celadonite; chl: chlorite; idd: iron oyxhydroxide; py: pyrite; sap: saponite; zeol: zeolite. The phases are listed in order of abundance.

140.75

138,97

Aphyric pillow basalt Aphyric massive basalt Vesicular aphyric massivc basalt Aphyric massive basalt

Rock Type

0.05

0.07 0.04 0.05 0.05 0.06

W/Roxy (open)

133.43

1028A-15X-7, 67-71cm, pc.9 (47°51.479'N 128°30.289'W) 1029A-25X-3, 99-105cm, pc. 14 (47°49.901"N 128°22.597'W) 1031A-6X-1, 46-49cm, pc.10 (47°53.400'N 128°22.597'W) 1032A-12R-2, 98-106cm, pc. 14 (47°46.773'N 128°07.34FW) 1032A-13R-3, 75-83cm, pc.12 (47°46.773'N 128°07.341'W) 1032A-14R-3, 104~108cm, pc.15 (47°46.773'N 128°07.341 'W) 0.4% 0.7%

1.8%

0.4%

0.4%

1.2%

0.5%

1.0%

Trace

Trace

1.5%

Sparsely phyric pillow basalt Moderately phyric pillow basalt Aphyric pillow basalt Aphyric pillow basalt Aphyric pillow basalt Aphyric pillow basalt 0.5%

Primary phenocrysts Plag. O1.

Rock Type

sap-cel-idd-pyhem sap-py

sap-carb-cel-iddPY sap-carb-cel-iddpy sap-carb-cel-iddph-hem

Pervasive

Analysis

8 8 12 12

7

Rim Core Rim Core

Pervasive

2 Rim < 1 Core 2 Rim

12

Secondary alteration Phases %

Trace sap-cel-idd-py-

2.2%

-

Px.

0.702620 0.702643 0.702615 0.702654

0.702575

0.702560 0.702484 0.702553

0.702624

87Sr/86Srn

7.95 7.79 7.75 7.45

7.44

7.80 6.56 7.10

7.49

6180

0.47 0.51 0.46 0.53

0.39

0.36 0.20 0.35

0.48

W/Rsr (open)

0.14 0.13 0.13 0.1 t

0.11

0.13 0.05 0.07

0.10

W/Roxy (open)

* Secondary phases are: carb: carbonate; cel: celadonite; hem: hematite; idd: iron oyxhydroxide; py: pyrite; sap: saponite; talc: talc; zeol: zeolite. The phases are listed in order of abundance.

313.20

303.59

292.65

41.76

220.79

Depth (mbsf)

Hole-core, interval, piece (Grid reference)

Table l(b). Summary of tithologicat units, primary and seconda~9, mineralogy, O'pe of analysis and 87Srff6Sr and 6180 isotopic ratios of samples from the Buried Basement (BB) transect.

*

173.78

1024C-1R-1, 8-15cm, pc.2 (47°54.531 'N 128°45.005'W) 1025C-4R-3, 36-44cm, pc.3A (47°53.247"N 128°38.919'W) 1025C-5R-1,137-140cm, pc, 14B (47°53.247'N 128°38.919'W) 1025C-5R-3, 29-37cm, pc.2 (47°53.247'N 128°38.919'W)

31.22

Depth (mbsf)

Hole-core, interval, piece (Grid reference)

Table l(a). Summary of the lithological units, primary and secondary mineralogy, type (ffanalysis and 87Srff6Sr and bmO isotopic ratios of samples from the Hydrothermal Transition ( HT) trans'ect

267.65

287.12

1026B-3R-t, 25-28cm, pc.3 (47°45.759'N 127°45.552'W)

1026B-5R-1, 5~56cm, pc.1 (47 45.759'N 127 45.552'W) 1027C-1R-1, 65 70cm, pc.lD (47 45.387'N 127 43.867'W) 1027C-1R-6, 12-17cm, pc.1 (47 45.387'N 127 43.867'W) 1027C-4R-1, 16-18cm, pc.3A (47 45.387'N 127 43.867'W) 1027C-4R-3, 76-81cm, pc.9 (47 45.387'N 127 43.867'W) 1027C-5R-5, 101 109cm, pc.9 (47 45.387'N 127 43.867'W) 1027C-5R-6, 3-8cm, pc.1 (47 45.387'N 127 43.867'W)

Dolerite-aphyric basalt maragin Moderately phyric pillow basalt Sparsely phyric pillow basalt Sparsely phyric pillow basalt Sparsely phyric pillow basalt

Basalt hyaloclastite breccia Aphyric pillow basalt Dolerite

Rock Type

1.2%

1.1%

1.0%

2.5%

9.0%

3.0%

Trace

0.8%

0.8%

0.8%

5.5%

-

28%

0.4%

Trace

Primary phenocrysts Plag. O1.

sap-cel-idd

sap-cel-py

sap-cel-py-iddzcol sap-py-carb-chl zeol-talc sap-chl-py-zeoltalc sap-cel-idd-hem

sap-carb-zeol

16 14

7

13 13 15

12

0.5 25 100 6 4 15

Secondary alteration Phases %

sap-py-carbhem-talc

-

Trace

1.2%

-

53%

0.2%

-

Px.

Rim Core

Pervasive

Rim Core Pervasive

Pervasive

Glass Basalt frag. Clay matix Rim Core Pervasive

Analysis

0.703422 0.703020

0.702642

0.702601 0.702602 0.702982

0.702577

0.702410 0.703056 0.706461 0.702533 0.702538 0.702662

S7Sr/86Srn

8.22 7.91

7.08

7.40 7.47 8.40

7.23

6.12 9.t2 19.26 7.60 6.70 8.63

b~SO

1.46 1.05

0.52

0.45 0.45 1.01

0.55

0.03 1.10 3.35 0.32 0.33 0.74

W/Rsr (open)

0,17 0.15

0.09

0.11 0.12 0.19

0.10

>

>

>

©

©

1.94

0.12 0.06 0.20

© © 0.02 0.24

W/Roxy (open)

* Secondary phases are: carb: carbonate; cel: celadonite; chl: chlorite; hem: hematite; idd: iron oyxhydroxide; py: pyrite; sap: saponite; talc: talc; zeol: zeolite. The phases are listed in order of abundance.

630.22

629.74

617.40

613.86

591.02

585.45

Depth (mbsf)

Hole-core, interval, piece (Grid reference)

Table l(c). Summary of" lithological units, primary and secondary mineralogy type o.1"analysis and 87Sr/XSSr and ~180 isotopic ratios of" samples.from the Rough Basement ( RB) transect.

106

A.G. HUNTER ET AL.

age from 1.46-1.95 Ma (Davis et al. 1997; Fig. l b, c), with the basement-sediment interface temperature increasing from 40-57°C (Davis et al. 1992; Davis et al. 1997), whilst the measured lithospheric heat flow is close to expected conductive values according to Davis et al. (1997) (Fig. 2a). Initial geochemical analyses indicate that the hydrothermal fluid contains a modified sea water component (Davis et al. 1997). As with the HT transect, igneous basement penetration varied from < 5m at Sites, 1208, 1029 and 1031, to ~ 5 0 m at Site 1032 (Fig. 2b; Davies et al. 1997).

represented by pillow basalts (found at all sites), massive basaltic flows (Holes 1024C, 1025C and 1027B), basaltic-hyaloclastite and basaltic breccia (Holes 1026B and 1027C) and a dolerite intrusive sheet (Hole 1027C). Recovery was poor in the rubbly pillow basalts (few percent), but improved significantly ( = 70%) within the more massive and coherent flow units. The following sections briefly describe each of the four rock types (Table 1), with a full description of petrological and textural variations, depth of penetration and recovery at each Hole available from the ODP Initial Reports volume 168 (Davis et al. 1997).

Rough Basement

P i l l o w basalts

(RB)

transect, sites

1026-1027

At ~100km east of the ridge axis, there is an abrupt change in basement topography from relatively smooth (BB transect), to a much rougher terrain (Fig. 2a). The Rough Basement (RB) transect consists of a series of parallel basement highs and troughs, with a local relief of 300-500m (Davis et al. 1992; Davis et al. 1997), produced by normal block faulting associated with variations in the magma supply rate at the time of crustal formation. Due to the irregular basement relief, the overlying sedimentary cover varies in thickness from a few tens of metres on basement highs, to several hundred metres in the troughs (Fig. 2a, b). Although the sedimentary cover is continuous within the RB transect, to the north of this area Mottl et al. (1998a, b) encountered exposed basement highs, some of which were surrounded by shimmering hydrothermal waters. From direct observations associated with geochemical analyses, Mottl et al. (1998a, b) have suggested that these highs enhance buoyancy-driven fluid flow between the crust and the ocean. Two sites were drilled in the RB transect; Site 1026 is located near the top of a basement high, while Site 1027 is located in the adjacent trough (Fig. 2a). The sites intersect crust of 3.49 Ma and 3.59 Ma, and are capped by sediments varying from 0.28 Ma to 1.68 Ma, respectively (Figs lc). Across the transect, the basement-sediment interface temperature varies from 61-64°C (Davis et al. 1997; Fig. 2a). This transect saw the greatest basement penetration, ranging from ~41 m at Site 1026, to ~ 5 0 m at Site 1027 (Fig. 2b; Davies et al. 1997).

Basement recovery and lithological units Across the ridge flank, four different emplacement modes were recognized in the basement,

Pillow basalts were recovered from all Sites, and although slight variations in the primary mineralogy and whole rock texture were noted, these did not change in a systematic manner. The majority of pillow basalts are aphyric (< 1% phenocrysts) to sparsely phyric (1-2% phenocrysts), and contain phenocrysts of plagioclase+olivine+pyroxene. In addition to the aphyric to sparsely phyric pillow basalts, minor amounts of moderately phyric (2-10% phenocrysts) plagioclase+olivine+pyroxene pillow basalts were also recovered at Sites 1026-1029. Texturally, the pillow basalts vary from glassy to hypohyaline and hypocrystalline, with vesicularity sparse to moderate (< 3.5 %). All of the pillow basalts have a typical depleted MORB signature (mg# = 0.70-0.58), with the compositional range attributed to varying amounts of low pressure fractional crystallization. M a s s i v e basalt f l o w s

Within Holes 1024C, 1025C and 1027B, the aphyric to moderately phyric plagioclase + olivine pyroxene basalts have a massive structure, with phenocrysts set in a cryptocrystalline to microcrystalline groundmass. As such, these rocks are believed to represent massive flow sequences, which are interbedded at points with occasional thin sedimentary horizons. Vesicularity in these basalts is generally sparse to moderate; however at Hole 1025C a progressive change in vesicularity occurs, increasing from 16mm and 5-15% from the top down to the middle of the flow units (Davis et al., 1997), before decreasing in size (2-1 mm), and abundance (5-< 1%) in the lower half of the flow units. This has been attributed to variations in the degassing and cooling rates in different parts of the flow (Davis et al. 1997). Geochemically, the massive basalts from

PETROLOGY OF LOW TEMPERATURE ALTERATION Holes 1024C and 1027B are compositionally similar to the pillow basalts, and can be related to the pillows by low pressure fractional crystallization. In contrast, the massive basalts from Hole 1025C are extremely evolved in composition (mg# <0.49; high TiO2 and Fe~_O3(total) wt%), and are indicative of ferrobasalts that have either formed from a late stage differentiated melt, or have been subject to a range of different physical conditions in the magma chamber compared with the depleted-MORB basalts, at the time of their formation (Davis et al. 1997).

Basaltic-hyaloclastite and basalt breccia Breccia was recovered from Holes 1026B and 1027B. In Hole 1026B, the breccia consists of angular to subangular basalt clasts and hyaloclastite shards, set in a matrix of granular saponite, carbonate and zeolites. The clasts are variably replaced by pale brown granular saponite :t: carbonate, and exhibit concentric alteration from the rims to the cores. The breccia from Hole 1027B consists of angular to subangular, aphyric basaltic clasts set in a muddy sedimentary matrix. The clasts vary in colour (attributed to differential amounts of hydrothermal alteration) and texture, and are believed to represent broken fragments from different levels within a cooling unit. Texturally and mineralogically, the basalt clasts in both breccias are indistinguishable from the aphyric to sparsely phyric basalts recovered from across the flank.

Dolerite sheet In Hole 1027C, a plagioclase-olivine-pyroxene dolerite sheet characterized by a holocrystalline, intergranular, subophitic to ophitic texture was recovered. Grain size varies from fine to medium throughout the unit, grading to a very fine grained, holocrystalline, aphyric basalt at its base; there is no systematic variation in modal abundance. The extent of alteration in the dolerite varies from slight to moderate (622%), and decreases towards the finer margins. Although the upper contact was not recovered, this unit was defined as a shallow intrusive body that may have been emplaced either on-axis, or during a later, separate off-axis stage of magmatism (isotopic analyses suggest the pillows and dolerite sheet originate from the same magmatic source, implying intrusion either on or adjacent to the axis for the sheet.)

107

Analytical techniques Twenty samples were selected from the HT, BB and RB samples for detailed petrological and mineral chemistry investigations. The samples exhibit a range of primary mineralogies and textures, as well as a range of secondary alteration types and phases. Based on the type of alteration recognized at the macroscopic and microscopic scales, the samples were subdivided into two groups consisting of: (i) samples with fresh ( < 2 % ) to moderate ( < 25%) pervasive alteration; and (ii) rim-core pairs, with the oxidation rims and non-oxidative cores separated prior to analysis (Descriptive terms for different amounts of alteration are based on the nomenclature used in the ODP Leg 168 Initial Reports volume, Davies et al. 1997)

Electron microprobe analyses. Chemical analyses of secondary phases were determined using a Cameca SX-50 electron microprobe (fitted with three wavelength dispersive spectrometers and a Link Analytical AN10/55S EDS system) at the University of Leeds, UK. The operating conditions used to analyse the secondary phases were: 15 kV accelerating voltage; 10 nA beam current; 5 #m beam diameter; with a counting time of a 10 s per element on the peak followed by a 10 s per element background. The beam current and beam diameter were decreased and increased respectively from normal conditions (i.e. 15 nA and 2-5 #m), in order to compensate for the unstable nature of the clays during analysis. All clay analyses were recalculated on a 22-oxygen formula-unit basis, with a selection of microprobe data listed in Table 2.

Stages of low temperature hydrothermal alteration. Low temperature alteration (< 150°C; (Honnorez 1981) of young oceanic crust is influenced by a number of variables relating to the basement rock (e.g. lithology, structure, permeability) and the alteration fluid (e.g. composition, temperature, fO2). As a result, the extent and type of alteration that can occur is extremely variable, both within and between different regions of the crust. Bass (1976) carried out one of the earliest low temperature hydrothermal alteration studies on oceanic basalts from the East Pacific Ocean (DSDP 34). From this and subsequent studies (see below), an alteration model has been proposed in which

108

A . G . HUNTER E T AL.

Table 2(a). Microprobe analyses for chlorite/smectite from the H T and RB transects Transect Hole Rock Type Location* Replacing Mineral~Colour

HT 1025C massive core mesostasis chl blue-green

RB 1027C dolerite core olivine chl blue-green

RB 1027C dolerite core plagioclase chl blue-green

RB 1027C dolerite core olivine chl blue-green

RB 1027C dolerite core olivine chl blue-green

RB 1027C dolerite core olivine chl blue-green

37.43 0.06 12.21 0.07 5.76 0.08 24.00 1.34 0.06 0.03 0.31 81.35

34.79 0.00 17.41 0.00 2.63 0.18 29.68 0.35 0.12 0.09 0.35 85.59

36.17 0.05 15.16 0.00 6.27 0.37 27.31 0.40 0.13 0.06 0.20 86.12

36.75 0.07 12.52 0.00 9.47 0.37 25.49 0.46 0.09 0.05 0.16 85.43

38.91 0.00 11.38 0.01 6.77 0.17 28.44 0.23 0.11 0.08 0.01 86.11

40.36 0.00 10.67 0.07 5.68 0.29 29.86 0.42 0.09 0.13 0.01 87.59

6.432 d.l. 2.472 < d.1. 0.840 0.000 6.144 0.240 0.024 0.000 0.048 16.248 0.880 6.984

5.170 < d.1. 3.036 < d.1. 0.330 0.022 6.556 0.066 0.044 0.022 0.044 15.268 0.952 6.908

5.434 < d.1. 2.684 < d.1. 0.792 0.044 6.116 0.066 0.044 0.022 0.022 15.202 0.885 6.952

5.654 < d.1. 2.266 < d.l. 1.210 0.044 5.852 0.066 0.022 0.000 0.022 15.180 0.829 7.106

5.830 < d.1. 2.024 < d.1. 0.858 0.022 6.358 0.044 0.022 0.022 0.000 15.180 0.881 7.238

5.918 < d.1. 1.848 < d.1. 0.704 0.044 6.534 0.066 0.022 0.022 0.000 15.180 0.903 7.282

SiO2 TiO2 A1203 Cr203 FeOT MnO MgO CaO Na~O K20

SO3 Total Si 4 + Ti 4+ A14 +

Cr 3+ Fe 2+ Mn 2+ Mg 2+ Ca 2+ Na + K+ S2Total rag# oct#

<

< d.l. = less than detection limit. *Location refers to location of alteration phase in rock sample. tMineral abbreviations:- chl: chlorite/smectite, mg# = Mg2+/[Fe 2+ + Mg 2+ + Mn2+]; oct#=Ti 4+ +A13+ +Fe 2+ + M n 2+ + M g 2+.

the stages of alteration c o r r e s p o n d with progressive burial of the oceanic crust a c c o m p a n i e d by changes in the c o m p o s i t i o n of the seawaterderived fluids, and by a successive shift in the conditions of alteration from open and oxidative, to restricted and reducing conditions. Such progressive patterns of alteration have been described a n d i n t e r p r e t e d by a n u m b e r o f workers at a wide range of locations, including: B6hlke et al. (1980), D S D P 46, N o r t h MidAtlantic; N a t l a n d & M a h o n n e y (1981), D S D P 60, Central West Pacific; H o n n o r e z et al. (1983), Hole 504B, Central East Pacific; Laverne and Vivier (1983), D S D P 70, Central East Pacific; Alt et al. (1986), Hole 504B, Central East Pacific; A d a m s o n & Richards (1990), O D P Leg 106, Mid-Atlantic; Alt et al. (1992), O D P Leg 129, Western Pacific; Alt (1993), O D P Leg 136,

Central Pacific; Laverne et al. (1996), O D P Leg 148, Central East Pacific; and Teagle et al. (1996), O D P Leg 148, Central East Pacific. Alteration p e t r o g r a p h y a n d mineralogy. T h e degree o f a l t e r a t i o n , defined as the m o d a l a b u n d a n c e of secondary minerals present in the rock, varies f r o m fresh ( < 2 % ) to m o d e r a t e (~<25%) across the eastern flank of the J d F R , although glassy rims on the pillow basalts are generally fresh ( < 2 % alteration) to pristine, whereas interstitial glass varies from fresh ( H T transition only) to completely altered. Basalt clasts f r o m the b a s a l t i c - h y a l o c l a s t i t e ( H o l e 1026B), exhibit slightly m o r e alteration (2540%) than other rock samples, although as with the pillow basalts, their glass rims are generally fresh ( < 0 . 5 % alteration). This contrasts with

8.030 0.000 0.110 0.000 3.652 0.000 1.122 0.088 0.044 1.408 0.066 14.498 0.235 4.774

Si4+ Ti 4+ AI 4+ Cr 3 + Fe 2+ Mn 2 + Mg 2+ Ca 2+ Na + K+ S2Total mg# oct#

6.996 0.000 0.396 0.000 5.214 0.022 1.430 0.132 0.088 0.902 0.044 15.224 0.215 6.666

40.14 0.08 1.88 0.04 35.86 0.19 5.54 0.68 0.24 4.10 0.28 89.03

1.140 0.008 0.128 0.000 1.172 0.004 0.196 0.036 0.004 0.148 0.008 2.852 0.143 1.380

33.35 0.36 3.20 0.03 41.01 0.17 3.88 0.99 0.06 3.42 0.24 86.73

RB 1027C pillow oxid halo fracture idd orange-red

1.156 0.008 0.132 0.000 1.132 0.000 0.216 0.036 0.004 0.140 0.004 2.828 0.160 1.356

34.65 0.24 3.37 0.04 40.48 0.02 4.36 1.00 0.04 3.28 0.23 87.71

RB 1027C pillow oxid halo vesicle idd orange

7.612 0.022 0.990 0.000 3.652 0.000 0.682 0.110 0.022 1.320 0.044 14.476 0.157 4.356

47.75 0.09 5.27 0.00 27.35 0.03 2.90 0.70 0.10 6.52 0.30 91.02

RB 1027C pillow oxid halo fracture cel green

7.546 0.000 0.176 0.000 5.016 0.022 0.814 0.088 0.022 1.320 0.000 15.026 0.140 5.852

43.43 0.00 0.87 0.00 34.51 0.08 3.17 0.53 0.04 5.96 0.06 88.65

RB 1027C pillow oxid halo vesicle cel-idd yellow

7.986 0.000 0.770 0.000 3.454 0.022 0.638 0.264 0.022 0.858 0.022 14.014 0.156 4.114

52.64 0.09 4.28 0.05 27.21 0.09 2.84 1.61 0.09 4.38 0.13 93.39

RB 1027C pillow oxid halo vesicle cel-nont green

*Location refers to location of secondary phase in rock sample. tMineral abbreviations:- cel: celadonite; idd: iron oyxhydroxide; nont: nontronite; sap: saponite. rag# = Mg2+/[Fe 2+ + Mg 2+ + Mn2+]; oct # = Ti 4+ + A13+ + Fe 2+ + Mn 2+ + Mg 2+.

51.85 0.03 0.58 0.05 28.18 0.00 4.87 0.49 0.16 7.08 0.57 93.88

SiO2 TiO2 A1203 Cr203 FeOT MnO MgO CaO Na20 K20 SO3 Total

Transect BB BB Hole 1032A 1032A Rock Type pillow pillow Location oxid halo oxid halo Replacing vesicle fracture Mineral t cel-nont cel-sap-idd Colour green yellow-green

Table 2(b). Microprobe analyses for oxidative secondary mineral phases in the B B and R B transects

7.854 0.022 0.638 0.000 3.674 0.000 0.748 0.110 0.022 1.364 0.022 14.454 0.169 4.444

50.18 0.14 3.41 0.02 28.07 0.00 3.22 0.71 0.06 6.8 0.16 92.77

RB 1027C pillow oxid halo vesicle cel-nont green

8.206 0.000 0.418 0.000 3.454 0.000 0.748 0.198 0.022 0.990 0.022 14.058 0.178 4.202

52.87 0.01 2.28 0.00 26.56 0.07 3.26 1.24 0.08 4.95 0.13 91.45

RB 1027C pillow oxid halo olivine cel-nont-sap green-brown

RB 1027C pillow oxid halo olivine cel-sap green-brown 51.9 0.36 6.55 0.11 24.92 0.01 2.51 1.12 0.13 5.32 0.25 93.18 7.832 0.044 1.166 0.022 3.146 0.000 0.572 0.176 0.044 1.034 0.022 14.014 0.154 3.762

RB 1027C pillow oxid halo vesicle cel-nont-sap green-brown 54.88 0.09 1.44 0.01 25.18 0.00 3.96 1.50 0.13 3.64 0.13 90.96 8.404 0.000 0.264 0.000 3.212 0.000 0.902 0.242 0.044 0.704 0.022 13.816 0.219 4.114

m

©

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110

A . G . HUNTER E T AL.

Table 2(c). Microprobe analyses for non-oxidative phases from the H T transect. Transect Hole Rock Type Location* Replacing Mineralt Colour

HT 1025C massive core olivine sap beige

HT 1025C massive core olivine sap grey-brown

HT 1025C massive core vesicle sap beige

HT 1025C massive core vesicle sap grey-brown

HT 1025C-5R1 massive core pyroxene talc colorless

HT 1025C-5R1 massive core pyroxene talc colorless

SO3 Total

51.01 0.03 3.74 0.03 17.47 0.10 17.01 1.14 0.42 0.17 0.52 91.64

46.34 0.07 3.91 0.01 16.31 0.06 15.42 1.50 0.21 0.09 0.91 84.83

49.78 0.00 4.04 0.01 17.15 0.05 16.51 0.92 0.91 0.24 0.48 90.10

44.76 0.08 3.53 0.01 16.58 0.08 14.37 1.48 0.18 0.16 0.98 82.21

60.28 0.27 0.97 0.00 2.04 0.08 28.33 0.14 0.17 0.02 0.15 92.45

59.17 1.01 0.95 0.00 5.52 0.11 26.98 0.11 0.20 0.02 0.09 94.16

Si4 + Ti 4 + A14+ Cr 3+ Fe 2+ Mn 2+ Mtg-2+ Ca 2+ Na + K+ S2Total Mg# Oct#

7.392 0.000 0.638 0.000 2.112 0.022 3.674 0.176 0.110 0.022 0.066 14.234 0.635 5.808

7.260 0.000 0.726 0.000 2.134 0.000 3.608 0.242 0.066 0.022 0.110 14.190 0.628 5.742

7.370 0.000 0.704 0.000 2.112 0.000 3.630 0.154 0.264 0.044 0.044 14.344 0.632 5.742

7.282 0.000 0.682 0.000 2.244 0.000 3.476 0.264 0.066 0.022 0.110 14.190 0.608 5.720

7.920 0.022 0.154 0.000 0.220 0.000 5.544 0.022 0.044 0.000 0.022 13.970 0.962 5.786

7.788 0.110 0.154 0.000 0.616 0.022 5.302 0.022 0.044 0.000 0.000 14.036 0.896 6.050

SiO2 TiO2 A1203 Cr203 FeOT MnO MgO CaO Na20

K20

*Location refers to location of alteration phase in rock sample. tMineral abbreviations:- sap: saponite; talc: talc. mg# = Mg2+/[Fe 2+ + Mg 2+ + Mn2+]; oct # = Ti 4+ + A13+ + Fe z+ + Mn 2+ + Mg 2+.

the basaltic-hyaloclastite matrix, which represents a complete secondary mineralogy, consisting of a m i x t u r e of s a p o n i t e + zeolites + carbonate, filling intraclast interstices. T h r o u g h o u t the eastern flank, alteration is present either as discrete a l t e r a t i o n haloes (parallel to rock margins and fracture surfaces, e.g. Fig. 3), a n d / o r as pervasive alteration. In both cases, the secondary phases either: (i) partially to completely replace olivine (with pyroxene and plagioclase rims also partially replaced in some samples); (ii) line/fill p r i m a r y void spaces and interstitial areas, i.e. vesicles, vugs and the mesostasis; and/or (iii) line/fill secondary fractures. The

main

alteration

phases

recorded

in

samples from all three h y d r o t h e r m a l transects include several phyllosilicates (e.g. saponite, celadonite, chlorite/smectite, talc), tectosilicates (e.g. quartz, zeolites), iron oxyhydroxides (e.g. hematite + the mineraloid iddingsite) a n d carbonate. The relative a b u n d a n c e of each alteration phase varies across the flank as well as within each transect, with the most a b u n d a n t being saponite, c a r b o n a t e , c e l a d o n i t e , i r o n oxyhydroxides, chlorite/smectite, and talc. Within and between samples f r o m the same and different holes, a range of crystal m o r p h o l o g i e s (e.g. fibrous, granular, crystalline), colour a n d o c c u r r e n c e (e.g. replacing different p r i m a r y phases; filling p r i m a r y pore spaces and secondary fractures), have been logged (Davis et aL 1997). A l t h o u g h the phyllosilicates were initially categorized as either saponite, celadonite, chlor-

6.534 0.000 0.902 0.000 4.554 0.000 2.728 0.154 0.132 0.154 0.000 15.158 0.375 7.282

Si4+ Ti 4+ AI4+ Cr 3+ Fe z+ Mn 2+ Mg 2+ Ca 2+ Na + K+ S2Total Mg# Oct#

7.238 0.000 0.198 0.022 2.860 0.022 3.234 0,330 0.088 0.066 0.220 14.300 0.531 6.116

46.93 0.04 1.13 0.21 22.13 0.15 14.09 1.95 0.29 0.38 1.86 89.17

BB 1032A pillow oxid halo olivine sap honey

7.194 0.022 0.462 0.000 2.684 0.022 3.454 0.396 0.066 0.022 0.110 14.410 0.563 6.182

47.97 0.12 2.58 0.02 21.5 0.20 15.48 2.51 0.19 0.15 0.90 91.61

BB 1032A pillow oxid halo olivine sap honey

7.590 0.000 0.462 0.000 1.298 0.000 4.510 0.198 0.044 0.022 0.022 14.146 0.777 5.808

52.03 0.02 2.7 0.00 10.66 0.01 20.78 1.32 0.13 0.14 0.29 88.07

BB 1032A pillow core olivine sap grey-brown

7.612 0.000 0.418 0.000 1.188 0.000 4.642 0.176 0.022 0.022 0.022 14.146 0.796 5.830

54.03 0.04 2.54 0.00 10.16 0.08 22.18 1.14 0.1 0.18 0.3 90.75

BB 1032A pillow core olivine sap grey-brown

7.238 0.000 0.880 0.000 1.452 0.000 4.114 0.286 0.066 0.022 0.088 14.168 0.739 5.566

51.47 0.03 5.37 0.00 12.24 0.07 19.59 1.92 0.21 0.18 0.83 91.91

BB 1032A pillow core vesicle sap grey-brown

7.656 0.000 0.352 0.000 1.100 0.022 4.686 0.154 0.044 0.022 0.044 14.080 0.810 5.808

54.56 0.02 2.16 0.00 9.41 0.10 22.37 1.05 0.12 0.17 0.47 90.42

BB 1032A pillow core vesicle sap grey-brown

*Location refers to location of alteration phase in rock sample. tMineral abbreviations:- cel: celadonite; idd: iron oyxhydroxide; sap: saponite. mg# = Mg2+/[Fe 2+ + Mg 2+ + Mn2+]; oct # = Ti 4+ + A13+ + Fe 2+ + Mn 2+ + Mg 2+.

39.05 0.03 4.57 0.01 32.51 0.05 10.92 0.86 0.40 0.73 0.03 89.16

BB 1032A pillow oxid halo fracture sap-idd honey

SiO2 TiO~ A1203 Cr2Os FeO-r MnO MgO CaO Na20 K20 SO3 Total

Transect Hole Rock Type Location* Replacing Mineral t Colour

Table 2(d). Microprobe analyses for non-oxidative phases from the BB and RB transects RB

RB

RB

RB

6.556 0.000 1.408 0.000 1.386 0.000 4.862 0.220 0.066 0.044 0.066 14.652 0.778 6.248

43.97 0.00 8.05 0.02 11.04 0.05 21.89 1.42 0.20 0.28 0.64 87.56 6.006 0.000 1.870 0.000 1.298 0.022 5.698 0.088 0.044 0.022 0.000 15.070 0.814 7.018

39.39 0.01 10.44 0.04 10.26 0.12 25.09 0.61 0.13 0,11 0.08 86.28 6.116 0.000 1.760 0.000 0.990 0.000 5.896 0.110 0.044 0.022 0.022 14.982 0.856 6.886

41.26 0.01 10.01 0.06 8.06 0.04 26.66 0.75 0.13 0.14 0.23 87.35

6.424 0.000 1.562 0.000 1.364 0.000 4.994 0.044 0.198 0.616 0.000 15.202 0.785 6.358

43.32 0.06 8.91 0.00 I 1.04 0.08 22.60 0.27 0.70 3.26 0.01 90.24

1027C 1027C 1027C 1027C dolerite dolerite dolerite dolerite core core core oxid halo olivine olivine olivine olivine sap sap sap sap-cel grey-brown grey-brown grey-brown grey-brown

RB

7.282 0.000 0.814 0.000 1.540 0.000 4.180 0.308 0.022 0.022 0.066 14.212 0.731 5.720

47.88 0.00 4.49 0.08 12.18 0.05 18.42 1.83 0.05 0.08 0.57 85.63

1027C pillow core vesicle sap grey-brown

0.000 1.999 0.001 0.000 2.000

0.00 99.95 0.05 0.00 100.00

0.00 99.89 0.05 0.00 99.94

HT 1025C massive core vein carb xtln

0.000 1.998 0.001 0.000 1.999

0.00 99.89 0.08 0.03 100.00

0.00 99.84 0.08 0.03 99.95

HT 1025C massive core vein carb sieved

0.174 1.762 0.023 0.040 1.999

7.41 88.9 1.35 2.34 100.00

7.40 88.81 1.35 2.34 99.90

BB 1032A pillow core olivine carb xtln

0.014 1.925 0.005 0.057 2.001

0.59 95.90 0.27 3.24 100.00

0.58 94.65 0.27 3.20 98.70

BB 1032A pillow core vesicle carb sieved

0.154 1.829 0.008 0.009 2.000

6.57 92.45 0.46 0.52 100.00

6.58 92.58 0.46 0.52 100.14

BB 1032A pillow core vesicle carb xtln

0.06 1.76 0.12 0.06 1.999

2.48 87.20 6.78 3.54 100.00

2.46 86.69 6.74 3.52 99.42

RB 1027C dolerite core olivine carb fibrous

0.021 1.898 0.059 0.022 2.000

0.89 94.44 3.42 1.25 I00.00

0.88 93.64 3.39 1.24 99.15

RB 1027C dolerite core spherule carb fibrous

0.000 2.000 0.000 0.000 2.000

0.00 100.00 0.00 0.00 100.00

0.00 99.53 0.00 0.00 99.53

RB 1027C dolerite core vein carb xtln

0.000 1.998 0.001 0.000 1.999

0.00 99.90 0.08 0.02 100.00

0.00 98.10 0.08 0.02 98.20

RB 1027C pillow core olivine carb xtln

0.171 1.799 0.012 0.018 2.000

7.28 90.96 0.72 1.04 I00.00

7.28 90.97 0.72 1.04 99.17

RB 1027C pillow core olivine carb xtln

0.163 1.811 0.014 0.011 1.999

6.95 91.55 0.84 0.66 I00.00

6.89 90.79 0.83 0.66 99.17

RB 1027C pillow core vesicle carb xtln

*Location refers to location of alteration phase in rock sample. +Mineral abbreviations:- carb: carbonate. Texture abbreviations: xtln: crystalline. Carbonate analyses normsed to 100%; cations recalculated as 6-fold formulae.

0.000 1.997 0.002 0.000 1.999

0.00 99.85 0.13 0.02 100.00

MgCO3n CaCO3n FeCO3n MnCO3n Total

Mg 2+ Ca-2+ Fe z+ Mn 2+ Total

0.00 98.91 0.13 0.12 99.06

HT 1025C massive core vein carb fibrous

MgCO3 CaCO3 FeCO3 MnCO3 Total

Transect Hole Rock Type Location* Replacing Mineralt Texture

Table 2(e), Microprobe analyses for carbonate Jhom the HT, BB and RB transects

PETROLOGY OF LOW T E M P E R A T U R E ALTERATION

113

,--.~ ~

/

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Chlorite, chloritesmectite mixes

Fe-oxyhydroxides, celadonite, Saponite, pyrite

Carbonate

1

2

3

4

BB, RB HT (pillow basalts), HT, BB, RB (all rock types) HT, BB, RB (all rock types)

HT (massive basalts); RB (dolerite)

Occurrence

Mesostasis; interstitial glass; olivine + plagioclase rims Mesostasis; interstital glass; vesicles; olivine. Mesostasis; interstital glass; vesicles; olivine Vesicles; fractures; olivine

Filling/Replacing

Main Associations

Along fractures; pervasive.

Pervasive

Oxidation haloes

Chlorite-smectite; celadonite; carbonate Saponite

Saponite

Along fractures and Saponite+Fegrain boundaries oxyhydroxides

Location

Oxidative; open; high temperature (chlorite: < 200°C*; chlorite-smectite: 80_50oc #) Oxidative; open; low temperature (25m7C°*). Non-oxidative ( oxidative); restricted; low temperature (23-70°C ~) Non-oxidative; restricted to closed; low temperature (35-70°C #)

Conditions of Formation

Temperatures taken from literature: §Lawrence & Drever, 1981; McMurtry et al., 1983; Staudigel & Hart, 1983; *Schiffmann & Fridleifsson, 1991 Stakes & O'Neil, 1982; #Teagle et al., 1996.

Alteration Phase

Alteration Stage

location, associations and conditions of forrnation

Table 3. Summary of the four stages of alteration recognized across the eastern flank qf the Juan de Fuca Ridge, noting variations in the diagnostic mineralogy, occurrence,

m

;~

PETROLOGY OF LOW TEMPERATURE ALTERATION

115

ite/smectite or talc petrographically, discrete mixtures between two or more alteration phases are apparent from XRD analyses (carried out during Leg 168, see Davies et al. 1997) and electron microprobe techniques (this study). These mixtures form a series of parallel discrete bands sequentially filling vesicles and fractures (indicative of the changing physiochemical conditions of alteration over time e.g. changing fO2, salinity, temperature, X~ui,0, as well as intimate mixtures characterized by a range in colours between the end-member phases. A full description of the type (e.g. haloes versus pervasive) and abundance of alteration, as well as variations in the morphology of the secondary mineral phases can be found in Davis et al. (1997).

tures and between grain boundaries. This resulted in patchy alteration of the mesostasis and interstitial glass, as well as the rimming of olivine and plagioclase phenocrysts by bluegreen chlorite and chlorite/smectite mixtures (identified by XRD and electron microprobe analyses, and hence forth referred to collectively as chlorite; Tables 2a & 3). Chlorite is characterized by a massive to fibrous texture and is commonly associated with saponite+zeolite (most commonly phillipsite, Davies et al. 1997)+iddingsite, forming either as intimate mixtures or as a complex of two or three consecutive but discrete bands of chlorite/ smectite + saponite+zeolite (Davies et al. 1997).

Sequence of low temperature alteration. Preliminary petrological investigations carried out in this study, indicate that a progressive sequence of alteration occurs within the eastern flank of the JdFR, with the location and extent of each alteration stage varying both within and between the three hydrothermal transects (Table 3). Across the eastern flank, the four stages of alteration have been defined according to:

Stage 2." iron oyxhydroxide + celadonite; open, oxidative alteration. The earliest widespread

(i) the secondary mineral assemblage (e.g. oxidative versus non-oxidative phases); (ii) variations in the style of alteration (e.g. core-rim v. pervasive alteration); and (iii) the relative timing of alteration (based on petrographic observations, i.e. sequence of formation). From this, the alteration stages have been subdivided into two early non-oxidative to oxidative stages, associated with open fluid circulation, followed by two non-oxidative stages, indicative of restricted or closed hydrothermal circulation. Stage 1: chlorite + chlorite/smectite; open, nonoxidative+oxidative alteration. The earliest

stage of alteration based on petrological evidence, is restricted in occurrence to coarser sections of the massive ferrobasalts (Hole 1025C) and the dolerite sheet (Hole 1027C). Petrological and mineralogical constraints suggest that this stage occurred soon after the rocks crystallized at or near the ridge axis, prior to sedimentation as well as before large volumes of oxidizing fluids had passed through the crust, cooling it to its ambient temperature. At this time, low rates of water-rock interaction between still warm crust and the alterations fluids would have occurred under open conditions, with the fluids passing unrestricted along frac-

stage of low temperature hydrothermal alteration in the eastern flank based on petrological observations, is represented by iron oxyhydroxides and celadonite, forming varicoloured alteration haloes. Experimental work on the stability of these phases (e.g. Garrels 1984; Donnelly et al. 1979; Lawrence & Drever 1981; Stakes & O'Neil 1982; Staudigel & Hart 1983), suggests that they would have formed in an open, oxidative environment, in an off-axis setting away from the main heat source, whilst the crust was either at or near the sea floor. Within the BB and RB transects, single oxidation haloes (e.g. iron oxyhydroxide; celadonite; or iron oxyhydroxide + celadonite) and double oxidation haloes (e.g. iron oxyhydroxide followed by a band of celadonite + iron oxyhydroxide) are abundant and very well defined (Fig. 3; Tables 2b & 3). This contrasts strongly with the HT transect, where examples of celadonit e + i r o n oxyhydroxide are very scarce, and when present, form narrow ( < 7 r a m ) , faint oxidation haloes. Iron-oxyhydroxides (e.g. hematite and the mineraloid iddingsite) are easily identified in hand specimen and thin section by their diagnostic bright orange to orange-red colour, and are characterized by either a massive, amorphous or sigmoidal texture, partially to completely filling fractures and vesicles. Celadonite (a dioctahedral K-rich illite-like mineral) meanwhile, is characterized by a distinctive bright to dark green colour, and varies texturally from amorphous, to granular or fibrous, partially to completely filling vesicles and fractures, and replacing patches of mesostasis. In general, celadonite is either associated with saponite (as intimate mixes or sequential fills), or follows iron oxyhydroxides as a discrete band in sequentially

116

A. G. HUNTER ET AL.

filled vesicles and fractures. Stage 3: saponite +pyrite; transitional oxidative to non-oxidative, restricted alteration. The first non-oxidative stage of alteration is marked by the formation of saponite (a trioctahedral smectite), closely associated with secondary sulphides (primarily pyrite), and is found in all three transects (Tables 2c, d & 3). Saponite is characterized by a beige to brown-grey colour, granular to amorphous texture, and generally replaces olivine, patches of mesostasis and interstitial glass, as well as partially to completely filling vesicles and fractures. It occurs as a solitary alteration phase, as well as being associated with a range of other phases, forming discrete bands in sequential fills (e.g. associated with celadonite+iddingsite, or carbonate) or intimate mixtures (e.g. mixed submicroscopically with celadonite, carbonate, pyrite, and/or zeolite). In a number of samples, the boundary between the oxidative and non-oxidative alteration zones is marked by a discrete band of secondary pyrite, separating the oxidative zone in which pore spaces have been filled by celadonite ± i r o n oxyhydroxide (+ olive-brown to honey brown saponite), from the pervasively altered core in which non-oxidative phases (e.g. grey-brown to beige saponite + pyrite ± talc +zeolites) are prevalent (Fig. 3), i.e. saponite forms a transitional phase between oxidative (saponite only) and non-oxidative (saponite + pyrite) stages of alteration. Variations in the colour and texture of saponite in the different alteration zones are suggestive of a change in composition, varying as the physicochemical conditions of hydrothermal alteration changed over time. Talc is limited to the replacement of olivine phenocrysts, and is closely associated with saponite (Tables 2c & 3). It occurs across the flank, but is most abundant in HT and BB transects. Stage 4." carbonate + zeolites + quartz, non-oxidative, restricted to closed alteration. The final stage of alteration is represented by the formation of carbonate and trace amounts of tectosilicates (e.g. zeolites, quartz; Tables 2e & 3). Carbonate is present in varying amounts in all three transects, filling fractures and vesicles, as well as replacing olivine phenocrysts. Texturally carbonate is diverse, exhibiting a range of morphologies including fibrous, granular and crystalline forms; both calcite and aragonite were identified across the flank by XRD (Davis et al. 1997).

XRD also identified several different zeolite species across the flank, with most common being phillipsite (Davis et al. 1997). Morphologically, the zeolites are fibrous to acicular, and are predominantly found replacing mesostasis in the basalts and/or forming fine rims in vesicles or round phenocrysts. M i c r o p r o b e analyses

Microprobe analyses have been carried out on a range of secondary phases from across the three hydrothermal transects, a summary of which is given in Table 2. When compared with published data from altered oceanic crust (Fig. 4, light grey fields), the alteration phases from the eastern flank of the JdFR exhibit systematic shifts in their mineral chemistries away from the expected compositions. For example, saponite and celadonite analysed in sections from the eastern flank of the JdFR, are characterized by high A1 contents and either low (saponite) or high (celadonite) interlayer cation totals (i.e. tetrahedral + octahedral totals; Fig. 4; Table 2b, d), compared with analyses from other altered oceanic crustal sections. The compositional ranges exhibited by the JdFR samples indicate that the clay analyses are not typical end member compositions produced by low temperature hydrothermal alteration of young oceanic crust, but represent mixtures of a number of alteration phases. In general, saponite from all three transects is characterized by higher mean Fe, A1, Ca and K contents but lower mean Si contents (Fig. 4; Table 2c, d), than typical Mg-rich saponites found in young oceanic basalts (e.g. Adamson & Richards 1990; Clauer & Chaudhuri 1995; Laverne et al. 1996). Celadonite from the BB and RB transects also differs slightly from the typical celadonite composition found in young oceanic crust (Fig. 4), characterized by higher A1 and Fe, and lower Mg abundances (Table 2b). Only the chlorite with interlayer cation totals of ~15, approximates to the expected end member composition (Fig. 4; Table 2a). In addition to the mineral specific groups, a mixed group of analyses with anomalously high FeOa-wt% can be recognized on Fig. 4 (i.e. group circled by the solid black line). The cause of this shift in mineral chemistry is discussed below. Discussion of mineral chemistry. Throughout the eastern flank of the JdFR, the secondary mineral assemblage varies on a regional and local scale. Petrological and microprobe analyses indicate that the main alteration phases (i.e. chlorite, iron

PETROLOGY OF LOW T E M P E R A T U R E A L T E R A T I O N

117

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Si/(Si+Al)

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45

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40 35 30

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65

%

Fig. 4. (a) mg# versus Si/(Si + AI) for secondary alteration mineral analyses from the HT (massive ferrobasalts), BB (pillow basalts) and RB transects (where RBb = pillow basalts and RBo = dolerite), illustrating the range in mineral compositions due to: mixing between different end member clay components; lithological influences; and changing physiochemical conditions of alteration. The analyses enclosed by the solid black line highlight the effects of interaction between iron oyxhydroxide and the different mineral phases. The light grey fields for celadonite, chlorite, saponite and talc represent the expected compositional ranges for each phase, taken from literature (incl. Deer e t al. 1983; Adamson & Richards 1990; Schiffman & Fridleifsson 1991; Alt 1993; Zierenberg e t al. 1995; Laverne e t al. 1996; Teagle e t al. 1996). rag# = Mg2+/(Mg 2+ + Mn + + Fe2+); (b) FeOT wt% versus SiO2 wt% for secondary alteration mineral analyses from the HT (massive ferrobasalts), BB (pillow basalts) and RB transects (where RB~ = pillow basalts and RBd -- dolerite), illustrating the range in mineral compositions due to: mixing between different end member clay components; lithological influences; and changing physiochemical conditions of alteration. Solid black line and grey fields as in Fig. 4a.

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oyxhydroxide, celadonite, saponite and carbonate) vary in appearance and composition. Although a number of different morphologies exist, the observed compositional variations do not correspond with systematic changes in either crystal structure or occurrence (i.e. what is being replaced/filled). Correlations can however be made between the composition of the alteration phase and its colour; the fabric and composition of rock in which it was analysed; and the interaction between fluids from different alteration stages (Fig. 4; Tables 2 & 3). Similar correlations have also been noted by several other workers (e.g. Honnorez et al. 1983; Alt et al. 1986; Adamson & Richards 1990; Laverne et al. 1996; Teagle et al. 1996), while studying low temperature alteration of upper oceanic crust from a number of locations. Correlation between colour and alteration mineral composition. The variations in colour in each of

the main alteration phases can be directly related to changes in cation concentrations (e.g. Laverne et al. 1996). For example, the saponite colours (e.g. grey-brown, olive-brown, honeybrown and beige) are produced by changes in the relative concentrations of FeOa-wt% and MgOwt%. In general, beige to grey-brown saponites characteristically contain moderate to high abundances of FeOTwt% and MgOwt%, whilst olive-brown to honey-brown saponite have low to very low FeOTwt% associated with moderate to high abundances of MgOwt%. This is illustrated on Fig. 4, where saponite from the BB transect forms two distinct groups characterized by differing mg# and SiO 2 wt% concentrations, according to their colour. Likewise, the range in colours exhibited by celadonite corresponds with different abundances of FeOa- wt%, MgO wt% and K20 wt%, altering from dark forest green (high Fe, low Mg, medium K), to bright emerald green (high K, medium to high Fe, low Mg) to pale yellow-green (medium Fe, and K, medium to low Mg). The intensity of the blue-green colour of chlorite, is correlated to the relative abundance of SiO2 wt% and in particular A1203 wt%, with increasing amounts of A1 strengthening the colour. As these secondary phases form from the alteration fluid, either directly by precipitation in primary and secondary void spaces, or indirectly by the replacement of magmatic phases, glass or mesostasis, any observed shift in mineral composition beyond that expected can be used to infer a change in fluid composition. Hence a change in the water/rock ratio and/or the physiochemical conditions of alteration can be

deduced (Clauer & Chaudhuri 1995). Effects of local textural and compositional variations on secondary mineral composition. Interac-

tion with the oceanic crust results in the composition of the alteration fluid changing. Thus, by changing the amount of water/rock interaction, and/or by varying the composition of the crust within the hydrothermal system, the fluid (and secondary mineral) composition will vary. Across the eastern flank, variations in composition (e.g. evolved ferrobasalts versus the more primitive tholeiitic pillow basalts) and texture of the basement rocks (e.g. pillow basalts versus the dolerite sheet) can be seen to have had an influential affect on the composition of the alteration minerals (Fig. 4; Table 2). For example, variations in the average composition of saponite mirrors the compositional shifts observed in the basement rocks, with saponite from the dolerite characterized by a higher average rag# (saponite average rag# = 0.81 + 0.04; dolerite average mg# = 0.67 4-0.02), than saponite from either the pillow basalts (saponite average mg#=0.65+0.16; whole rock average rag# = 0.634-0.02) or the massive ferrobasalts (saponite average rag# = 0.626-1-0.01; whole rock average rag# = 0.49 + 0.05) (Table 2c, d). As well as the geochemical influences, variations in the local crystallinity, grain size and permeability of the basement have also influenced alteration mineral compositions, by controlling the efficiency with which the fluid has interacted with the crust (i.e. the water/rock ratios). Across the flank, samples characterized by a coarser average grain size, higher crystallinity and/or higher abundance of void space (primary and/or secondary) have generally undergone more water/rock interaction that adjacent samples (Table 1), and contain alteration phases systematically enriched in MgO, CaO, K20 and SiO2 wt%, but depleted in FeO-r, MnO and TiO2 wt% (Table 2), compared to secondary minerals in finer grained and/or less permeable, but compositionally comparable rocks. Mixing between two or more alteration fluids and/ or secondary phases. Finally, variations in

alteration mineral chemistry can also be partially attributed to mixing between different end member clay phases, which are themselves indicative of the changing physiochemical conditions of alteration. Petrological observations indicate that in many samples, a selection of different alteration phases are commonly found either as consecutive bands in vesicles and fractures and/or forming intimate mixtures.

PETROLOGY OF LOW TEMPERATURE ALTERATION These intimate mixtures consist of two or more phases that have intergrown at a scale finer than can be analysed by electron microscopy, thus appearing as intermediate compositions between different clay components, e.g. the data circled by the solid black line on Fig. 4. In addition, the two groups of identically coloured saponites analysed in the dolerite on Fig. 4, are compositionally distinct due the presence/absence of previous alteration phases in the host rock (Table 2c, d). Saponite analysed in chlorite-bearing zones is characterized by higher A1203 wt%, rag# and octahedral number, but lower SiO2 wt% than the saponite from otherwise unaltered zones (Fig. 4). Thus, submicroscopic intergrowths of saponite and chlorite have shifted the bulk composition of saponite away from the expected compositional field, with these intergrowths produced by mixing between the alteration fluid responsible for saponite formation and previously formed chlorite in the host rock. Contamination of the alteration fluid with previously formed secondary phases has also resulted in the formation of fine-scale intergrowths between saponite + celadonite + iron oyxhydroxide, talc + celadonite, celadonite + iron oyxhydroxide, and iron oyxhydroxide + chlorite (Fig. 4; Table 2). Evolutionary sequence o f low temperature hydrothermal alteration

Using the variations in mineral chemistry discussed above, in association with petrological observations, a detailed summary of hydrothermal alteration within the eastern flank of the JdFR can be produced. Changes in the mineral assemblage as well as mineral chemistry, have been used to interpret variations in the physiochemical conditions of alteration (e.g. oxygen potential; fluid and basement temperature; fluid chemistry etc.) in the basement, as it moved away from the buoyant ridge axis and was progressively buried by sedimentation (+offaxis magmatic events). Stage 1: magmatic-deuteric, non-oxidative+ oxidative alteration. The earliest stage of alteration in the eastern flank is restricted to the massive ferrobasalts (HT transect) and dolerite sheet (RB transect), and is represented by the formation of chlorite and chlorite/smectite mixes. An estimate of the temperature and conditions of formation for these phases can be made based on the observed range in compositions (this study), associated with previously published experimental work on the

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stability of chlorite (Bird et al. 1984; Schiffman & Fridleifsson 1991) and chlorite/smectite (Teagle et al. 1996). From this, it is probable that the chlorite formed at <200°C (Schiffman & Fridleifsson 1991), at or near to the ridge axis in a non-oxidative environment (as the majority of iron in chlorite is normally in a ferrous state; only small amounts of ferric iron are occasionally found in chlorite, Deer et al. 1982), with open unrestricted fluid circulation occurring at a time when the crust was still warm, prior to it being cooled down to ambient temperatures by large volumes of oxidizing fluids passing through the crust. As the system cooled and the water/rock ratio increased, petrological observations indicate that chlorite was superseded by the formation of chlorite/smectite at c. 80-50°C (Teagle et al. 1996) in a similar to slightly more restricted, and more oxidative environment. Compositionally, the chlorite found in the eastern flank could not have formed from unmodified seawater (e.g. MgO and A1203 wt% in chlorite are too high). This implies that if the fluid involved in this stage was originally seawater, it must either have undergone significant water/rock interaction prior to chlorite formation in order to modify the fluid chemistry as required, or represent a mixture of seawater and deeper, enriched crustal fluids. Studies on similar occurrences of early formed chlorite and chlorite/smectite in oceanic settings propose that these phases are produced by interaction between the crust and a variety of different fluids derived from either magmaticdeuteric, metasomatic and/or hydrothermal sources. From work on Holes 504B (DSDP Leg 70) and 896A (ODP Leg 148), Laverne (1987) and Laverne et al. (1996) suggested that chlorite formed by a reaction between hot oceanic crust and late stage magmatic fluids, released during the last stages of crystallization, whereas Alt et al. (1986) suggested the chlorite in Hole 504B formed by reactions between seawater and hot crust during initial cooling periods. During ODP Leg 168, the direct measurement of up-flowing fluids in the RB transect suggests that chlorite formed by the interaction of the crust with a mixture of seawater and compositionally modified fluid from depth. Although a similar upflow zone was not directly observed in the HT transect, the presence of high temperature (< 600°C) CaNaTi-rich pyroxenes along some fractures in the massive ferrobasalts indicates that late-stage magmatic-deuteric derived fluids were present in this crustal region, allowing a model for

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chlorite formation to be proposed, in which the oceanic crust interacted with seawater and a compositionally modified hydrothermal and/or late stage magmatic fluid sourced from depth. Therefore, this stage of alteration represents initial interaction between newly formed, warm crust with seawater and a late stage magmatic and/or hydrothermal fluid, occurring prior to the onset of wide spread low temperature hydrothermal alteration.

have been compared with published isotopic and mineralogical studies on celadonite, in order to estimate a potential temperature of formation. From this, temperatures of between 25-47°C (Stakes & O'Neil 1982; McMurtry et al. 1983; B6hlke et al. 1984) can be proposed for celadonite formation in the JdFR. The lack of significant quantities of celadonite within the HT transect (in particular Sites 1023 and 1024) may be due to:

Stage 2." open, oxidative, hydrothermal alteration. The earliest, widespread stage of low temperature (~< 150°C) hydrothermal alteration in the eastern flank occurred once the crust had moved off-axis, away from the underlying heat-source whilst the crust was either still at or near the seafloor. The lack of a significant sedimentary cover at this time, allowed open hydrothermal circulation under oxidizing conditions, associated with a steady flux of unmodified to slightly modified seawater. As the water/rock ratios increased (Table 1), interaction between low temperature seawater derived fluids and the oceanic crust by diffusive exchange between the hydrothermal fluids flowing through cracks and fractures, and the pore fluids held in the crust, resulted in the oxidative breakdown of primary iron-bearing phases (e.g. olivine and primary sulphides). This led to the formation of iron oxyhydroxides (e.g. hematite, iddingsite) at the outermost edge of oxidation haloes and pillow rims, as well as along fractures. Continued water/rock interaction between the crust and seawater associated with progressive modification of the fluid composition, subsequently resulted in the formation of celadonite (a bright green dioctahedral K-rich illite-like mineral; Fig. 3; Table 2b). The presence of celadonite in the oxidative haloes indicates that the physiochemical conditions of the system must have been further modified after iron oyxhydroxide formation. The generally higher abundance of ferric rather than ferrous iron in celadonite (Adamson & Richards 1990) compared with the iron oxyhydroxides (Table 2b), is characteristic of a slight decrease in the fO2 of the system. Secondly, before celadonite can form, the alteration fluids must be enriched in Fe, Si and alkali elements relative to the average abundances of these elements in pure seawater, i.e. sufficient water/rock interaction is required to alter the fluid composition before celadonite can form (Adamson & Richards 1990). Although no O- or Sr-isotope analyses have yet been carried out on celadonite in this study, the mineral chemistry analyses that have been performed,

(i) chemical constraints, with insufficient water/rock preventing the composition of the seawater-derived fluid to be modified sufficiently to allow celadonite to form; (ii) temperature constraints, with the measured basement-sediment interface temperatures of < 25°C at Sites 1023 and 1024 (Davis et al. 1997, Fig. 2a) too cold for this phase to form; or perhaps (iii) sampling constraints, with celadonite forming at greater depths in the basement than was sampled at Sites 1023 and 1024. This contrasts with the BB and RB transects (and eastern margin of the HT transect) in which celadonite is abundant, and is found at all depths in the crust (from < 1-50 m), and where water-rock interaction is higher and the measured basement-sediment interface temperature is also higher, ranging between 38 ° to 64°C (Davis et al. 1997). Stage 3: restricted, non-oxidative hydrothermal alteration. Continued burial of the basement due to sedimentation and/or periods of off-axis magmatism associated with progressive formation of alteration minerals in the crust, restricted fluid flow and changed the system from an open to slightly restricted, oxidative environment, to a more restricted or closed, non-oxidative environment. At the same time, continued waterrock interaction modified the composition of the seawater derived fluid further, increasing the relative abundance of Fe, Mg, Si and alkali elements in solution, by the breakdown of magmatic phases and glass (see Table 1 and discussion below). These changes in the physiochemical conditions of alteration resulted in the pervasive formation of saponite (a trioctahedral MgFe-smectite) throughout the crust, with the shift from oxidative to non-oxidative alteration transitional, forming mixtures of celadonite and saponite (as evident petrographically and in the mineral chemistry), prior to the onset of true non-oxidative alteration and the formation of saponite+pyrite (Fig. 4; Table 2b, d). As with celadonite in the previous alteration

PETROLOGY OF LOW TEMPERATURE ALTERATION stage, potential temperatures of formation for saponite have been calculated from a combination of mineral chemistry (this study) and published mineral and isotopic analyses. From this, the compositional range of saponites recovered from the eastern flank appear to have formed over a range of temperatures, from ~<70°C (Stakes & O'Neil 1982; Staudigel & Hart 1983), to 23°C (Lawrence & Drever 1981; McMurtry et al. 1983), continuing down to perhaps as low as 0°C (B6hlke et al. 1984). Thus, the wide thermochemical conditions under which saponite is stable explains why this phase is found across virtually all of the eastern flank of the JdFR (except Site 1023 due to insufficient water-rock interaction), with the formation and range in observed saponite mineral chemistry predominantly controlled by the fluid composition, the occurrence of non-oxidative alteration conditions and the fluid-basement temperature. S t a g e 4: late restrictive, non-oxidative alteration.

The formation of M g - F e rich clays (e.g. saponite) during low temperature alteration of the crust, preferentially removes Mg cations from the fluid, whilst the progressive breakdown of magmatic phases such as plagioclase, increases the relative abundance of Ca in the fluid, as the water/rock ratio increases (Mottl et al. 1983). As the alteration continues, the oxidation potential of the system decreases, whilst the pH of the fluid increases, encroaching the stability ranges required for the final stage of alteration represented by carbonate (+ zeolite + authigenic quartz) to occur (Mottl et al. 1983; Alt & Honnorez 1984). Based on initial petrological observations, two discrete stages of carbonate formation were recognized across the flank. The earliest carbonate, which formed either just after or synchronously with saponite, is limited to lining/filling fractures and is found in all three transects. This was followed by the later carbonate stage which post-dates saponite formation, filling vesicles as well as replacing olivine phenocrysts and patches of mesostasis; this carbonate is only found in the BB and RB transects. Preliminary microprobe analyses confirm the existence of two discrete stages of carbonate formation (Table 2e). The first stage is characterized by a pure Ca-carbonate (containing > 99 wt% CaCO3), whilst in the second stage, the composition varies, consisting of a C a M g + ( M n + Fe) carbonate, with the relative abundance of Fe and Mn increasing whilst Ca decreases with distance from the ridge axis. The presence of Fe and Mn in the later carbonate implies that the physiochemical conditions of

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alteration must be sufficiently reducing in order to allow these cations to be removed from the basement, and mobilized in the fluid (Alt & Honnorez 1984; Clauer & Chaudhuri 1995). The increasing amounts of Fe and Mn in the carbonate with distance from the ridge axis, indicates that as the crust moved away from the spreading centre, the oxidative potential of the system progressively decreased. Teagle et al. (1996) recognized two similar stages of carbonate formation in Holes 504B and 896A, and used oxygen isotope analyses to calculate temperatures of formation of ~<35°C for the Ca-carbonate and ~<70°C for the Ca-Mgcarbonate. Previous work on low temperature alteration of oceanic crust (Andrews 1977; Hart et al. 1994; Laverne et al. 1996) has suggested that carbonate can only form in relatively old crust (i.e. ~<2.5 Ma), limited by the time required to sufficiently modify the composition, pH and oxidation potential of the seawater-derived alteration fluid in order to reach the stability conditions of carbonate. However, within the eastern flank of the JdFR, varying amounts of carbonate are present in all three transects, forming in crust varying in age from 0.99 Ma (Site 1024) to 3.59 Ma (Site 1027). This implies that, presuming the alteration fluid in the eastern flank of the JdFR is derived from seawater, the rate at which the composition of the fluid is being modified must be higher than normal, allowing carbonate to form in relatively young crust. The actual cause of this higher modification rate has yet to be quantified, but may be influenced by a variety of factors including local and regional variations in the lithology of the basement and overlying sediments, the thickness and lateral continuity of these units, the primary and secondary permeability of the basement and sediments; the occurrence, size and connectivity of faults, and most importantly, the basementsediment interface temperature. Thus, within the eastern flank of the JdFR, changes in the extent and duration of waterrock interaction, as well as changes in the fluid chemistry have been recorded by the alteration mineral assemblage in the basement, with the diagnostic phases of each stage recording the progressive change in physiochemical conditions of alteration. In general, alteration in the eastern flank commenced under open oxidative conditions, characterized by a low water-rock ratio between fresh to slightly modified seawater and shallow buried basement. This resulted in the breakdown of magmatic phases, associated with the formation of a range of secondary oxidative phases (e.g. iron oyxhydroxide, celadonite) with-

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in discrete to well defined alteration haloes (Fig. 3). Over time, the physiochemical conditions of alteration progressively changed, with a less oxidative environment characterized by restricted circulation of more highly modified seawater-derived fluid, resulting in the formation of a non-oxidative mineral assemblage including saponite, pyrite and carbonate.

Effects of low temperature alteration on young versus old oceanic crust From the discussion above, the sequence of events, processes involved and resultant effects of low temperature hydrothermal alteration within the eastern flank of the JdFR are directly comparable to the findings of many other studies on young (< 10 Ma) oceanic crust (e.g. Bass 1976; Andrews 1977; B6hlke et al. 1980; Natland & Mahonney 1981; Honnorez et al. 1983; Laverne and Vivier 1983; Alt et al. 1986, 1996; Adamson & Richards 1990; Laverne et al. 1996; Teagle et al. 1996), making the findings of this study widely applicable. Recent work on low temperature hydrothermal alteration of old (> 10Ma) in situ oceanic crust on ODP Legs 129 (Alt et al. 1992; 95 Ma crust, Hawaiian Arch, Central Pacific) and 136 (Alt 1993; 158165 Ma crust, east of the Mariana Trench, western Pacific), have shown that the progressive sequence of low temperature hydrothermal alteration generally found in young oceanic crust and across the eastern flank of the JdFR, can also be recognized in older upper crustal sequences. In general, the amount of bulk rock alteration in older crustal sequences was found to be quantitatively comparable to that in younger crust (i.e. < 30% bulk rock alteration, Alt 1993). This implies that after the initial period of low temperature alteration in the first few million years after basement formation, the oceanic crust does not undergo a significant amount of further alteration as it ages. By using Rb-Sr dating techniques on celadonite and saponite veins in the upper crust from several DSDP Sites, Hart & Staudigel (1986) have suggested that a time limit of ~<10Ma after basement formation can be placed on the main stages of oxidative (iron oxyhydroxides + celadonite) and non-oxidative (saponite + pyrite) alteration in the oceanic crust. After this time, the crust remains in a relatively steady state, with alteration limited to the formation of minor amounts of carbonate (i.e. 2-3%, Alt et al. 1986, 1992), with the increasing size and abundance of latestage calcite veins in ageing upper crustal

sequences, the main difference between young and old crust (Staudigel et al. 1981; Alt et al. 1992; Alt 1993). Where older oceanic crust was found to exhibit higher than expected levels of low temperature alteration, additional physiochemical factors were found to have influenced the rate (and extent) of alteration. For example at Site 801 (ODP Leg 129), the moderate to high (30-80%) amounts of alteration recorded in the off-axis eruptive alkali basalt unit were attributed by Alt et al. (1992) to slightly higher ambient crustal temperatures produced either directly from the off-axis eruptive units or from an adjacent active magma chamber. These higher ambient crustal temperatures in turn enhanced the temperature at which alteration occurred, and increased the convection rate of the fluids through the crust and overlying sediments, leading to higher water-rock ratios and hence higher rates of alteration. Accelerated rates of alteration produced by slightly higher than normal ambient crustal temperatures may be used to explain the early occurrence of carbonate in the young (0.993.59Ma) crust across the eastern flank of the JdFR. Rapid burial of the still warm crust by variable thicknesses of sediment will have insulated the crust, enhancing both the rate at which water-rock exchange occurred and the rate at which low temperature alteration took place, reaching the physiochemical conditions required for carbonate formation within a relatively short time after basement formation. Therefore, petrologically, young and old oceanic crust exhibit the same extent and type of low temperature hydrothermal alteration, with the bulk of this occurring in the first 10 Ma after basement formation (Hart & Staudigel 1986; Alt et al. 1992; Alt 1993). Studies on anomalously altered young and old upper oceanic crust (e.g. the rapid rate of alteration in the relatively young crust across the eastern flank of the JdFR (this study), and the high alteration levels in the alkali basalts at Site 801, Alt et al. 1992), indicate that although low temperature hydrothermal alteration is controlled by a number of physiochemical factors (e.g. basement-sediment lithology, permeability, fluid composition), the most influential of these is the basement-sediment temperature.

Conclusions By combining petrological and petrochemical analyses carried out on the alteration phases in a range of stratigraphically well constrained samples from the eastern flank of the JdFR, a

PETROLOGY OF LOW TEMPERATURE ALTERATION detailed representation of low temperature hydrothermal alteration has been established. This study has shown that in agreement with previous work on both young and old oceanic crust, four stages of low temperature alteration can be recognized in the eastern flank of the JdFR. The different stages of alteration represent varying amounts of water-rock interaction in the crust, with this increasing with depth of burial at any one location in the flank, as well as with increasing distance from the ridge axis across the flank. As the basement has moved away from the ridge axis and become progressively buried by sediment, the physiochemical conditions of low temperature alteration have undergone systematic changes. Adjacent to the ridge axis, unmodified to slightly modified seawater derived fluids were able to flow freely through the basement, under oxidative conditions, with the hydrothermal regime characterized by a low water-rock ratio. Continued water-rock interaction associated with progressive sedimentation of the basement, restricted fluid-flow and recharge through the basement, modify the composition of the seawater-derived fluid as well as causing the physiochemical conditions of the system to change from an oxidative to non-oxidative environment. These changes in the hydrothermal regime not only affected the location of alteration but also changed the diagnostic secondary mineral assemblage that formed, from chlorite and chlorite/smectite in the earliest stages of alteration, to iron oxyhydroxides and celadonite under oxidative conditions, associated with open fluid circulation through the basement, to saponite 4-pyrite under more restricted, non-oxidative conditions, ending with the formation of carbon a t e + z e o l i t e s + q u a r t z by the interaction of strongly modified seawater derived fluids with the basement. Thus, throughout the eastern flank of the JdFR, continuous interaction between variably modified low temperature seawater derived fluids and the basement, has resulted in a progressive change in the basement secondary mineral assemblage, recording the shift from open, oxidative alteration to restricted or closed non-oxidation conditions of alteration. This research was supported by a Royal Society Dorothy Hodgkin Research Fellowship (502008. K507) to AGH. Thanks go to E. Condliffe (University of Leeds) for help and advice on the microprobe, and to the technicians onboard the JOIDES Resolution who made ODP Leg 168 a very enjoyable cruise. The author would like to thank J. Alt, J. Honnorez and an anonymous reviewer for their constructive corn-

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Contributions to Mineralogy and Petrology, 73, 341-364. CANDE, S. C. & KENT, D. V. 1992. A new geomagnetic polarity time scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research, 97, 13917-13951. CLAUER, N. & CHAUDHURI, S. 1995. Clays in crustal environments: isotope dating and tracing. 1st edition, Springer-Verlag, Berlin. DAVIS, E. E. & CURRIE, R. G. 1993. Geophysical observations of the northern Juan de Fuca Ridge system: lessons in seafloor spreading. Canadian Journal of Earth Sciences, 30, 278-300. --, CHAPMAN,D. S., MOTTL, M. J., et al. 1992. FlankFlux: an experiment to study the nature of hydrothermal circulation in young oceanic crust. Canadian Journal of Earth Sciences, 29, 925 952. ,FlSHER, A. T., FmTIq, J. V., et al., (eds) 1997. Proceedings of the Ocean Drilling Program, Initial Reports, 168. DEER, W. A., HOWIE, R. A. & ZUSSMAN,J. 1982. An Introduction to the Rock Forming Minerals'. 13th edition, Longman Group Ltd, England. DONNELLY,T. W., PRITCHARD,R. A., EMMERMANN,R. & PUTCHETT,H. 1979. The ageing of oceanic crust: synthesis of the mineralogical and chemical results of Deep Sea Drilling Project, Legs 51 through 53. In: DONNELLY, T. W., FRANCHETEAU,J., BRYAN, W., ROBINSON,P., FLOWER, M., SALISBURY,M. et al. (eds) Initial Reports of the Deep Sea Drilling Project, 51-53, part 2, 1563-1577. FISHER, A. T., BECKER, K., NARASIMHAN, T. N., LANGSETH, M. G. 8¢ MOTTL, M. J. 1990. Passive, off-axis convection through the southern flank of the Costa Rica rift. Journal of Geophysical Research, 95, 9343-9370. GARRELS, R. M. 1984. Montmorillonite/Illite stability diagrams. Clays and Clay Minerals, 32, 161-166. HART, S. R. & STAUDIGEL,H. 1986. Ocean crust vein mineral deposition: Rb/Sr ages, U-Th-Pb geochemistry, and duration of circulation at DSDP Sites 261,462 and 516. Geochimica Cosmochimica Acta, 50, 2751-2761. - - , BLUSZTAJN,J., DICK, H. J. B. & LAWRENCE,J. R. 1994. Fluid circulation in the oceanic crust: contrast between volcanic and plutonic regimes. Journal of Geophysical Research, 99, 3163-3174. HONNOREZ, J. 1981. The ageing of the oceanic crust at low temperature. In: EMILIANI, C. (ed.) The Sea, vol. VII." The oceanic lithosphere. Wiley, New York. - - , LAVERNE, C., HUBBERTEN, H. W., EMMERMANN, R. & MUEHLENBACHS,K. 1983. Alteration processes in layer 2 basalts from Deep Sea Drilling Project Hole 504B, Costa Rica Rift. In: CANN, J. R., LANGSETH,M. G., HONNOREZ,J., VON HERZEN, R. P., WHITE, S. M., et al. (eds) Initial Reports of the Deep Sea Drilling Project, 69, 509546. LANGSETH, M. G., BECKER, K., VON HERZEN, R. P. & SCHULTHEISS,P. 1992. Heat and fluid flux through sediment on the western flank of the Mid-Atlantic Ridge: a hydrogeological study of North Pond. Geophysical Research Letters, 19, 517-520.

LAVERNE, C. 1987. Unusual occurrences of aegerineaugite, fassite and melanite in oceanic basalts (DSDP Hole 504B). Lithos, 20, 135-151. & VIVIER, G. 1983. Petrographical and chemical study of basement from the Galapagos Spreading Centre, Leg 70. In: HONNOREZ, J., VON HERZEN, R. P., et al. (eds) Initial Reports of the Deep Sea Drilling Project, 70, 375-390. - - , BELAROUCHI, A., & HONNOREZ, J. 1996. Alteration mineralogy and chemistry of the upper oceanic crust from Hole 896A, Costa Rica Rift. In: ALT, J. C., KINOSHITA, H., STOKKING,L. B. & MICHAEL, P. J. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 148, 151-170. LAWRENCE, J. R. & DREVER, J. I. 1981. Evidence for cold water circulation at DSDP Site 395: isotopes and chemistry of alteration products. Journal of Geophysical Research, 86, 5125-5133. MCMURTRV, G. M., WANG, C. H. & YEn, H. W. 1983. Chemical and isotopic investigations on the origin of clay minerals from the Galapagos hydrothermal mounds field. Geochimica Cosmochimica Acta, 47, 475-490. MOTTL, M. J., ANDERSON, R. N., JENKINS, W. J. & LAWRENCE, J. L. 1983. Chemistry of waters sampled from basaltic basement in DSDP Holes 501, 504B and 505B. Initial Reports of the Deep Sea Drilling Project, 69, 475-484. , WHEAT, C. G. & DAVIS, E. E. 1998a. Hydrothermal venting through outcrops on 3.5 Ma-old crust, eastern flank of the Juan de Fuca Ridge. Journal of Geophysical Research, in press. , --, BAKER, E., et al. 1998b. Warm springs discovered on 3.5 Ma-old crust, eastern flank of the Juan de Fuca Ridge. Geology, 26, 51-54. NATLAND,J. H. & MAHONNEY,J. J. 1981. Alteration in igneous rocks at Deep Sea Drilling Project Sites 458 and 459, Mariana Fore-arc Region: Relationship to basement structure. In: HUSSONG, D. M., UYEDA, S. et al. (eds) Initial Reports of the Deep Sea Drilling Project, 60, 769-788. SCHIFFMAN, P. 8¢ FRIDLEIFSSON, G. O. 1991. The smectite to chlorite transition in drillhole NJ-15, Nesjavellir geothermal field, Iceland: XRD, BSE, and electron microprobe investigations. Journal of Metamorphic Geology, 9, 679-696. STAKES, D. S. & O'NEIL, J. R. 1982. Mineralogy and stable isotope geochemistry of hydrothermally altered oceanic rocks. Earth and Planetary Science Letters, 57, 285-304. STAUDIGEL, H. 8¢ HART, S. R. 1983. Alteration of basaltic glass: Mechanisms and significance for the oceanic crust-seawater budget. Geochimica Cosmochimica Acta, 47, 337-350. - ~¢ DELANEY, J. R. 1981. Alteration of the oceanic crust: processes and timing. Earth and Planetary Science Letters, 52, 311-327. TEAGLE, D. A. H., ALT, J. C., BACH, W., HALL1DAY,A. N. & ERZINGER, J. 1996. Alteration of upper ocean crust in a ridge-flank hydrothermal upflow zone: mineral, chemical, and isotopic constraints from Hole 896A. In: ALT, J. C., KINOSHITA, H., STOKKING, L. B. & MICHAEL, P. J. (eds) Proceedings of the Ocean Drilling Program, Scientific

PETROLOGY OF LOW TEMPERATURE ALTERATION

Results, 148, 119-150. THOMPSON, R. E., DAVIS, E. E. & BURD, B. J. 1995. Hydrothermal venting and geothermal heating in Cascadia Basin. Journal of Geophysical Research, 100, 6121-6142. WHEAT, C. G. & MOTTL, M. J. 1994. Hydrothermal circulation, Juan de Fuca Ridge eastern flank:

125

factors controlling basement water composition.

Journal of Geophysical Research, 99, 3067-3080. ZIERENBERG, R. A., SCHIFFMAN, P., JONASSON,I. R., TOSDAL, R., PICKTHORN,W. & MCCLAIN, J. 1995. Alteration of basalt hyaloclastite at the off-axis Sea Cliff hydrothermal field, Gorda Ridge. Chemical Geology, 126, 77-99.

The structure and controls on fluid-rock interactions in ocean ridge hydrothermal systems: constraints from the Troodos ophiolite M. J. B I C K L E 1, D. A. H. T E A G L E 2, J. B E Y N O N 1 & H. J. C H A P M A N 1

1Department of Earth Sciences, University of Cambridge, Cambridge CB2 3EQ, UK 2Department o f Geological Sciences, University o f Michigan, Ann Arbor, M I 48109-1063, USA

Abstract: Rb-Sr isotopic compositions of rocks and minerals from the Troodos ophiolite have been analysed to constrain the rate limiting mechanisms that control fluid-solid exchange, the extent to which recharge fluids were channelled and the thermal evolution of oceanic hydrothermal systems. Systematic regional sampling has confirmed the Sr-isotopic alteration profile suggested by Bickle & Teagle (1992). This has been previously interpreted as a consequence of recharge fluids percolating down through the extrusive series with kinetically limited fluid-rock interaction but altering the underlying ~1 km section of sheeted dykes with near equilibrium fluid-rock Sr-isotopic exchange. Detailed Sr-isotopic profiles reported here across structural heterogeneities such as pillows, dyke margins and faults have failed to show isotopic gradients related to channelling of fluid. Rather the fluid flow in the recharge was pervasive and the degree of alteration is largely controlled by the extent of mineral recrystallization, primarily by albitization of plagioclase, and to a lesser extent by recrystallization of pyroxene. Hydrothermal recharge should cool the sheeted dykes much faster than the progression of the Sr-isotope alteration through the crust and rocks at greenschist facies (~250°C) or higher temperatures are predicted only to be found within a few hundred metres of the basal boundary layer of the system. The alteration pattern on Troodos with a zone a kilometre or more in thickness altered at greenschist facies temperatures and with significant Sr-isotope exchange is thus not easily explained.

The circulation of seawater through the ocean crust is a ubiquitous phenomena associated with the production of new oceanic crust at all types of oceanic spreading ridges. Although there have been major advances in the description and understanding of mid-ocean ridge vent systems since their discovery (Corliss et al. 1979), the magnitude, mechanisms, and geochemical consequences of hydrothermal circulation remains poorly constrained and controversial. In particular, the global hydrothermal fluid flux remains highly uncertain. Estimates based upon global chemical budgets of elements such as Mg and Sr (see review by Elderfield & Schultz 1996) imply cumulative fluxes that exceed, by up to an order of magnitude, those calculated from thermal constraints (Fig. 1; Fehn et al. 1983; Sleep 1983, 1991; Morton & Sleep 1985). Even if the flux estimates based on thermal modelling are increased by up to a factor of two to allow for higher estimates of the latent heat and thickness of oceanic crust (Hess 1992; White et al. 1992), reconciliation of the oceanic Sr-isotope budget with the thermal constraints, requires an average effluent temperatures of ~200°C. This conflicts

with the observational evidence that fluids are either heated to ~350°C or circulate at much lower temperatures ( < 100°C, B6hlke et al. 1984; Hess et al. 1991; Mottl & Wheat 1994) at shallower depths which seems to have limited impact on Sr-isotope budgets (Palmer & Edmond 1989). The discrepancy between the geochemical and thermal estimates of hydrothermal fluxes might be a result of inadequacies in either our understanding of the structure of oceanic h y d r o t h e r m a l systems or in other aspects of global geochemical cycles. Resolution of the problem is of primary importance for understanding of the controls on the composition of the Earth's oceans and atmosphere. Most existing models of ocean floor hydrothermal systems are constrained by observations on the chemistry of the hydrothermal effluents and the characteristics of the chimneys, vent structures and hydrothermal mounds (Tivey et al. 1995; Von D a m m 1995) or are based on physical models of hydrothermal circulation in the oceanic crust that either assume a permeability structure for the oceanic crust or use the few available constraints such as depth to the

BICKLE,M. J., TEAGLE,D. A. H., BEYNON,J. • CHAPMAN,H. J. 1998. The structure and controls on fluid-rock interactions in ocean ridge hydrothermal systems: constraints from the Troodos ophiolite In." MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the GeologicalRecord, Geological Society, London, Special Publications, 148, 127-152

127

128

M.J. BICKLE E T A L . Thermal Budget HydrothermalFlux Estimates

20

Heat loss all \\ ~ hydrothermal

15 Global Flux 10

13

\

kg.a -1

k

N

10-

Flux per unit area 7

Geochemical Budget 6 Estimates ~7Sr/86Sr

[ll

10 -2 kg.m

N

\

Mg*

\

133 He

[23 Li [4]

Probable hydrothermal heat loss i

00

i

,

•

2'00 3'00 Vent Fluid Temperature °C

i

•

0

400

Fig. 1. Estimates of hydrothermal flux from heat available in oceanic crust 7.1 km thick, thermal properties given in Table 1, assuming all heat lost by hydrothermal circulation (dashed line) and more plausibly ~50% efficiency (Sleep 1991) (solid line). The ocean crust is presumed to be cooled to 0°C down to 2.5 km and 860°C between 2.5 and 7.1 km. Global heat loss rate is calculated assuming seafloor spreading rate of 3.45km2a-l of Parsons (1982). Estimates from geochemical balance in the oceans after [1] Palmer & Edmond (1989), [2] Jenkins et al. (1978), [3] assumes river input of Berner & Berner (1987) but uncertainty in low temperature loss of Mg may reduce this estimate, [4] Chan et al. (1992).

magma chamber and effluent fluid temperatures (Fehn et al. 1983; Morton & Sleep 1985). However, models of the structure of the oceanic hydrothermal systems which fit the surface observations are not unique, despite the critical information on the temperature and depth of fluid-rock reaction zones provided by study of the chemistry of the effluent fluids. A complementary approach to the study of oceanic hydrothermal systems is the careful documentation of hydrothermal alteration preserved in in situ ocean crust and in ophiolites. Rocks and secondary minerals from these environments provide a cumulative record of fluid-rock interaction and fluid pathways which is not available from surface observations of active systems. In this paper we determine the geometry of the recharge fluid flow in ocean floor hydrothermal systems and the nature of the kinetic controls on fluid-rock geochemical exchange, by mapping the strontium isotopic alteration of the Troodos ophiolite and by examining the spatial and mineralogical controls on the degree of strontium-isotopic exchange. The work was undertaken to test and extend the isotopic tracer transport model for the Troodos ophiolite proposed by Bickle & Teagle (1992). The main conclusion is that the kinetics of fluid-solid strontium isotope exchange are largely controlled by the degree of recrystallization of the original igneous minerals. However, current

models of the structure of the hydrothermal systems with recharge fluxes large enough to cause the strontium-isotopic alteration do not predict high enough temperatures to form the > 1 km of greenschist facies alteration assemblages observed in the lower parts of the recharge zone.

The Troodos ophiolite The Troodos Massif, Cyprus is one of the best preserved and most widely studied ophiolites (Gass 1968; Moores & Vine 1971; Robertson & Xenophontos 1993). The extensive sheeted dyke complex provides unequivocal evidence for an origin in some form of oceanic spreading margin. Although the compositions of lavas from the Troodos ophiolite indicate formation in a supra-subduction environment (Miyashiro 1973; Rautenschlein et al. 1985), the igneous stratigraphy of the Troodos ophiolite is apparently indistinguishable from that of in situ ocean crust. The very limited post-emplacement deformation and metamorphism, render the superb 3D outcrops of the Troodos ophiolite an appropriate location for studies of oceanic hydrothermal circulation. Easily accessible, extensive outcrops of upper oceanic crustal rocks enable detailed descriptions and intensive sampling strategies to be implemented that would be prohibitively expensive and probably not technically feasible on existing oceanic crust.

CONSTRAINTS FROM THE TROODOS OPHIOLITE sea floor

~,

<

: Axis

0

:

129 Axis

0

,22

sheeted [ I I I N

depth km

IJ

Gabbros

2.5

Recharge on ult zone

dykes / / / / N High flux ~

2

t

discharge

Nth

Magma chamber~ <

2.5 -

ary

-

layer

. Axis 0

<

Axis

:

2.5

&

2.5

Fig. 2. Possible geometries for recharge and discharge paths of oceanic ridge hydrothermal circulation. (a) Uniform recharge, concentrated discharge. Note basal, high-temperature boundary layer may migrate laterally through gabbros as magma chamber crystallizes, (b) recharge and discharge both concentrated after Cann & Strens (1982), (c) recharge flux decreases downwards after Bowers & Taylor (1985); and (d) hydrothermal flow path is time-varying to flush anhydrite from crust after Sleep (1991).

Structure of oceanic hydrothermal systems: thermal and geochemical constraints Knowledge of fluid-flow paths, the extent of fluid-rock interaction and the porosity/permeability structure of the ocean crust are essential to predicting the impact of oceanic hydrothermal systems on the geochemistry of the oceans. Numerical models of hydrothermal circulation have assumed porous media flow through hot, spreading oceanic crust with a uniform permeability overlying an impermeable layer or with permeability decreasing exponentially downwards (Sleep & Wolery 1978; Fehn et al. 1983). These models suggest that the markedly nonlinear variation of the specific heat of water with temperature causes the hydrothermal systems to comprise broad regions of recharge and more focused zones of upwelling. The numerical models do not fully simulate the point source discharge of consistently high temperature fluid observed at black smoker sites. Cann & Strens (1982) and Cann et al. (1985) pointed out that the thermal constraint imposed by the fluxes of high temperature fluid emitted at black smoker sites requires that the fluid is heated in a narrow boundary layer of high permeability situated within tens of metres of a magma chamber that supplies the heat by

convection. The mapping of the narrow zones of the highly altered quartz-chlorite and epidosite rocks on the gabbro-sheeted dyke contact and in narrow distributed zones within the sheeted dykes in the Troodos ophiolite (Richardson et al. 1987; Schiffman & Smith 1988; Teagle 1993; Bettison-Varga et al. 1995) provides field evidence for such high-permeability boundary layers situated close to the top of magma chambers and for concentrated zones of discharge, presumably the feeders for black smoker complexes. The ~5 to 10% volume loss caused by formation of epidosite or quartz-chlorite rocks from diabase, provides a mechanism for the transient tbrmation of the very high porosity, high permeability zones required to concentrate the flow into the narrow boundary layers and discharge zones (Teagle 1993). The structure of the high temperature boundary layers and concentrated discharge zones provide no direct constraints on the structure of the recharge zones within the oceanic crust and a wide range of flow paths have been suggested (Fig. 2). Cann & Strens (1982), Cann et al. (1985) and Lowell et al. (1995) have suggested that cold down-welling fluids are channelled by high permeability structures (faults/detachments) to the base of the sheeted dyke complex and heated to black smoker temperatures by the conduction of heat across a narrow boundary

130

M. J. BICKLE E T A L . 87

90 Ma 0.703

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Sr/ Sr

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* Diabasedyke • CY4 x Faultrocks & zeolites o Published analyses

Fig. 3. 87sr/a6sr ratios, corrected to 90 Ma, through the Troodos crust showing previously published compilation by Bickle & Teagle (1992) as open symbols and new data as filled symbols. Depths estimated as in Bickle & Teagle (1992). 87Sr/S6Sr of fresh glass after Rautenschlein et al. (1985) and late Cretaceous seawater after Jones et al. (1994).

layer between the upper crust and lower convecting and crystallizing magma chamber. Other models (e.g. Bowers & Taylor 1985) have presumed decreasing fluid penetration with depth to simulate decreasing water/rock ratios, a model which is consistent with flow in a crust with a marked decrease in permeability with increasing depth (Rosenberg e t al. 1993). It is possible that only part of the flow is heated to ~350°C and much fluid is emitted from the seafloor as diffuse flow at intermediate temperatures and without forming the spectacular black smokers that have attracted so much attention. The thermal structure and flow field in the recharge zone are the principal controls on the extent of fluid-rock interaction for many of the important geochemical tracers within the hydrothermal systems. Observations on the chemistry of the vent fluids and the pattern and degree of alteration in old oceanic crust or ophiolites can be used to constrain the structure of the recharge part of the hydrothermal flow. Bickle & Teagle (1992) compiled an 87Sr/86Sr profile for the Troodos ophiolite and, integrated

with detailed mapping of the alteration, used this to constrain the geometry of fluid circulation paths, assess the magnitude of the timeintegrated fluid flux as well as constrain the extent of fluid-solid exchange for strontium isotopes in different parts of the hydrothermal system. The S7Sr/S6Sr profile of the Troodos crust (Fig. 3) exhibits substantial scatter in the pillow lavas spanning the complete range between unaltered basalt (~0.7035, Rautenschlein e t al. 1985) and contemporaneous seawater (0.7074, Jones et al. 1994). This contrasts with a restricted range of strontium isotopic compositions (87Sr/86Sr=0.7047-0.7059) for all rock types from the underlying ~ 1 km comprising the sheeted dyke complex and uppermost plutonic rocks. Rocks below the uppermost plutonics yield primary igneous SVSr/86Sr ratios. An important observation is that the quartzchlorite and epidosite zones, the probable pathways of concentrated high temperature fluid upwelling, have identical 87Sr/ 86 Sr to background diabase dykes. Since the highly altered epidosite and quartz-chlorite rocks were probably in

131

CONSTRAINTS FROM THE TROODOS OPHIOLITE Transition Sheeted Dykes Zone

Lavas 87

)

86

Sr/ Sr ~,'~" Seawater ~ ~I ~ 0.707 [ " "- ,, Fluid composition I

Minimum 87Sr/86Sr penetration on Troodos

/

0.706 L

0.705

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rock composition <

0.704

Rock & fluid composition

>1<

Kinetic Zone N D = 1 I

0

I

l

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I

2

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Fig. 4. Strontium isotope transport model proposed by Bickle & Teagle (1992). Seawater passes down through lavas with kinetically limited fluid solid exchange such that fluid and average rock are not in Sr-isotopic equilibrium. Within sheeted dykes, fluid and rock approach Sr-isotopic equilibrium. Model was drawn for timeintegrated flux of 3 x 107 kg m-- and the dimensionless Damk6hler Number = 1 for first-order kinetics (scaled for a distance of the time-integrated fluid flux). This causes fluid 87Sr/86Sr to vary from ~0.7045 to 0.7055 at the base of lavas over the life of the system. This time-varying signal is mapped into the sheeted dyke complex. Note in real oceanic crust the transition between kineticallylimited exchange and fluid-solid equilibrium is likely to be gradual and time-varying, precluding a sharp step in isotopic composition across the transition zone.

strontium isotopic equilibrium with the hydrothermal fluid, the implication is that much of the sheeted dyke complex in the recharge zone on the Troodos ophiolite was also altered to strontium isotopic equilibrium with the recharge fluid. The recharge fluid was set to the observed intermediate SVSr/S6Sr ratios (0.7047-0.7059) during passage through the overlying pillow lavas in which alteration was kinetically limited (Fig. 4). Bickle & Teagle (1992) made three significant inferences about the relationship between alteration and hydrothermal circulation on Troodos: (1) The uniformity of 87Sr/86Sr in the diabases of the sheeted dykes and the quartz-chlorite and epidosite discharge zones (~0.7054) indicates that strontium isotopic exchange took place during the high temperature phase of hydrothermal circulation and that the 87Sr/86Sr profile reflects alteration in the recharge zone of an evolving hydrothennal system, (2) the absence of very elevated 87Sr/86Sr from any rocks in the deeper parts of the hydrothermal system precludes significant channelling of recharge fluids along permeable fault zones and (3) the alteration of a ~1 kin thick sheeted dyke zone with near fluid-rock equilibrium allows calculation of the time-integrated fluid flux which reached the high-tempera-

ture boundary layer at the base of the system and was presumably discharged as high temperature black smoker fluid. Given a solid/fluid partition coefficient of Sr of ~10 by mass, Bickle and Teagle (1992) calculated the minimum time-integrated flux as 3 x l 0 7 k g m -2, a flux which is high compared with thermal estimates for the high T flux (maximum 107 kg m -z) but comparable with that based on oceanic mass balance of strontiumisotopes (3.5×107kgm -2, Palmer & Edmond 1989; Fig. 1). It should be noted that the strontium-isotopic systematics on Troodos im1Y that the ophiolite was discharging fluid with Sr/S6Sr ratios significantly elevated above that of the basalt. This contrasts with most active black smoker systems which discharge fluid with a strontium isotopic ratios less than 0.0015 above that of the underlying crust. Bickle and Teagle (1992) modelled strontiumisotope transport through the Troodos ophiolite using a one dimensional model with uniformly down-welling flow. The model considered the Troodos crust as two separate reaction regimes; an upper part comprising the pillow lavas in which fluid-solid exchange is kinetically controlled, and a lower region with near-equilibrium fluid-solid exchange (Fig. 4). The time-integrated fluid flux necessary to alter the ~1 km thick sheeted dyke section is calculated fi'om the relationship between the velocity of the altera-

132

M.J. BICKLE E T AL.

Table 1. Definition of variables and constants Variable

Definition

value

units

Cp Cs, Cf h L Kd Pe t Tm V Wo

Specific heat of basalt Concentrations of tracer in solid and fluid Distance Latent heat of fusion of basalt Solid/fluid partition coefficient by mass Thermal Peclet number, Pe = Woh/~ Time Magma temperature at ridge Velocity of alteration front Darcy fluid flux Thermal diffusivity Solid density Fluid density

1.3

kJ kg #gg-l m kJ k g 1

Ps, pf

tion 'front' (IQ, the partition of the tracer between solid and fluid by mass (Kd=Cs/Cf) where C~ and Cf are the concentrations of the tracer in the solid and fluid respectively and p~ and pf are the density of solid and liquid respectively, and the Darcy fluid flux (W o) (Table 1, Bickle & McKenzie 1987). It is given by:

Wo v - psK~/ p~f

(1)

such that if the alteration front is transported distance h (h = Vt where t is time), the timeintegrated fluid flux ( W o t ) is given by: W o t = h p~Kd . Pf

(2)

If the alteration takes places with fluid-solid exchange which is rapid compared with the transport velocity, V, the alteration will propagate as a relatively sharp 'front' across which the tracer concentration or isotopic composition rises from that of the unaltered basalt to that of the infiltrating fluid. Such fronts may be broadened over limited distances by diffusion and flow-path related dispersion and much more substantially by kinetically limited fluid-solid exchange. Bickle & Teagle (1992) show that for a simple first-order kinetic model, if the alteration in the overlying pillow lavas takes place sufficiently sluggishly, the fluid entering the top of the sheeted dykes may have its 87Sr/86Sr ratio composition buffered to a relatively narrow range between ~0.7045 and 0.7055 over the time necessary to drive the Sr-isotopic alteration front 1 km across the underlying sheeted dykes. The calculated time-integrated fluid flux is a m i n i m u m because the alteration front has passed entirely through the sheeted dyke zone

600

1250 10-6 3 x 103 103

s °C ms ms 1 ms 2 kg m 3 kg m 3

and into the high flux zones now denoted by epidosite and quartz-chlorite rocks. In this paper we investigate two important premises of the model: (1) that alteration in the sheeted dykes was to relatively uniform 87Sr/86Sr ratios; and (2) the nature of the kinetic rate law appropriate for strontium isotope exchange in the upper part of the system. The first premise, of uniform 87Sr/86Sr ratios within the sheeted dyke complex, has been tested by analysing samples collected over a much wider area. The second, the nature of the kinetic control, has been investigated by searching for gradients in 87Sr/86Sr ratio which might relate to transport away from fluid conduits and also by examining the relationship between mineralogical recrystallization and strontium isotopic alteration.

Analytical methods Sample localities are shown on Fig. 5. Rock samples were collected by hammering or drilling 5 to 8 cm long cores. Sr isotopic compositions were analysed in Cambridge on both the single collector VG54E mass spectrometer and the T40 VG Sector 54 multicollector mass spectrometer in dynamic mode. Repeat analyses of NBS987 gave a mean of 0.710241+82 on the VG54E and 0.710261+20 (2o- errors) on the T40 VG Sector 54 over the periods of the analyses. Rb and Sr concentrations were analysed by isotope dilution. Analytical, chemical processing methods and isotopic spikes were similar to those used by Bickle et al. (1988). Sr blanks were less than 1 ng and Rb blanks less than 0.3 ng. Rb/Sr ratios reproduce to better than 1% relative on samples

CONSTRAINTS FROM THE TROODOS OPHIOLITE

/--

/

f f ~ ~f~ ; ; 5¢,~,','2o~ 2~ o X ; o o o J ~ o2~ U : II ° ° I lavas

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.

~. . . . . . . . . . .

~c~ ~/ ")~( - g ~ ~ K M . . . . . . . . . 4

Gabbro

= I

I" / i; ".( ;7o c> <~ Z <~ " gK.LIKOU q )>~77 ? ? "

~ ~ fa.J , , /

.

I=

, PP89 Sample locality

1 / / /

@ = =_.~A /', A-OLYMPUS a A\. . . . . . . . q;

133

?

.

.

?

.

?

.

.

~

?.

-

]

7

°

(BLA

/4,,

m akheras f %'--;'Q" ' Monastery"~ "A ' 0 Kilometres

",~ 4

Fig. 5. Simplified map of northern part of Troodos ophiolite after Wilson & Ingham (1959), Bear (1960) and Carr & Bear (1960) showing sample localities and main geological units. Samples prefixed by MM collected adjacent to the Makheras Monastery. with more than 0.3 ppm Rb. 87Sr/86Sr ratios used below have been corrected to 90 Ma, the age of the ophiolite (defined by plagiogranites dated at 91.6+1.4Ma by U-Pb analyses on zircons, Mukasa & Ludden 1987) using the analysed Rb/Sr ratios.

Strontium isotopic composition of the basal group and the sheeted dyke complex recharge zone The transport modelling undertaken by Bickle & Teagle (1992) is based upon the inference that the recharge must have been relatively homogeneous and that alteration took place with near equilibrium fluid-rock Sr-isotope exchange in the sheeted dyke complex. These assertions were founded upon the limited, yet elevated, range of 87Sr/86Sr ratios (~0.7047-0.7059) exhibited by rocks of all alteration styles analysed from the sheeted dykes and uppermost plutonic complexes. The sampling of the sheeted dyke complex and uppermost plutonic rocks by Bickle and Teagle (1992) was strongly biased towards regions proximal (~1 km) to zones of epidositization and sulphide mineralization, both phenomena indicative of concentrated hydrothermal discharge (Richardson et al. 1987; Kawahata & Scott 1990; Teagle 1993). The major epidosite zones recognized to date

have a spacing of ,-~5-10 km. There was also a deficit of analyses of igneous textured diabase with actinolite + calcic plagioclase found in the deeper parts of the sheeted dyke complex (Baragar et al. 1989). In this study, these sampling deficits have been redressed in order to completely characterize the recharge zone by sampling the basal group and sheeted dyke complex systematically across the ,-~30km strike width of the ophiolite, supplemented by samples from the upper 630 m of the CY-4 drill hole (Baragar et al. 1989). The new analyses provide information on both the heterogeneity of recharge fluids distal to discharge zones and the extent to which primary STSr/S6Sr is retained by igneous minerals indicating significant fluid-rock disequilibrium in the lower part of the sheeted dyke complex. Samples were collected from within the basal group to the sheeted dyke-gabbro transition along road-cuts following north-south flowing rivers spaced ~ 5 k m apart on a west-east transect from Nikos to several kilometres northeast of the Makheras Monastery (Fig. 5). Detailed sampling was also made north-south along the Phterykhoudhi Potamos, from clinopyroxene-bearing diabase that crop out in roadcuts near the Makheras Monastery, and from the CY-4 borehole. The distribution of samples is shown in Fig. 5 and sample descriptions, locations and 87 Sr/ 86 Sr are presented in Table 2.

M. J. B I C K L E ET AL.

134

Table 2. Strontium isotopic compositions of samples from the regional study of the sheeted dyke complex and basal group Sample number

Depth (m)

GR

Sample description

87Sr/86Sr

2sigma

Sr ppm

Rb 87Sr/86Sr ppm (t = 90 Ma)

Samples ~omthe CCSP-CY-4 drillhole (Palekhori) 27.10 43.55 104.55 118.05 165.30 167.65 178.85 182.00 196.90 234.50 273.44 278.10 281.32 338.10 341.81 377.56 407.40 425.6O 441.75 447.20 460.70 491.05 557.70 581.80 609.30 623.50

1897 1914 1975 1988 2035 2038 2O49 2052 2067 2105 2143 2148 2151 2208 2212 2248 2277 2296 2312 2317 2331 2361 2428 2452 2479 2494

50852-386389

0.704662 0.704971 0.704538 Diabase: Fresh plag rare 0.704189 Cpx completely altered 0.704156 Act 30-50%, 0.704718 0.704119 Chl 0-5%, Ep 0 2% 0.704223 0,704863 0.705079 0.704446 0.704341 0.704112 0.704345 0.704482 0.703937 Diabase: Fresh plag 0.703700 common, 0.704386 some fresh Cpx, 0.704128 Act 30-50%, 0.704906 no Chl or Ep 0.704191 0.704941 0.704715 0.704615 0.705134 0.704386

52 28 42 54 22 28 24 24 24 30 26 22 22 20 36 26 36 48 34 22 34 22 30 34 60 26

107.0 75.2 112.6 89.2 101.5 97.7 107.1 105.5 77.9 146.7 74.0 73.1 103.3 98.5 90.2 67.2 74.3 71,0 97.6 69.5 78.0 126.9 106.7 80.0 96.2 46.5

0.5 0.4 0.2 0.5 0.2 0.4 0.4 0.3 0.5 0.4 0.5 0.4 0.4 0.6 0.2 0.1 0.2 0.2 0.4 0.3 0.2 0.7 0.2 0.4 1.0 0.2

0.704645 0.704950 0.704530 0.704170 0.704148 0.704704 0.704104 0.704214 0.704849 0.705079 0.704422 0.704323 0.704098 0.704323 0.704474 0.703933 0.703692 0.704376 0.704112 0.704890 0.704181 0.704920 0.704707 0.704598 0.705095 0.704367

Diabase-Act Diabase- Act Diabase- Act Diabase- Act, (Chl) Diabase- Act Diabase- Act, (Ep)

0.704811 0.705046 0.705371 0.704889 0.704642 0.704570

20 30 22 22 20 24

90.2 139.0 118.1 86.7 109.6 97.4

0.6 1.6 1.0 0.3 0.6 0.4

0.704784 0.705003 0.705338 0.704874 0.704622 0.704553

50740 387525 50760-387600 50770-387630

Diabase- Act, Chl Diabase- Act, Chl Diabase- Act, Chl

0.705127 0.705370 0.705487

56 24 38

93.1 74.9 49.6

1.8 1.1 0.2

0.705055 0.705316 0.705474

2439 2280 1988 1610

48405-387390 48415 387540 48420-387825 48515-388015

Diabase- Act, Chl Epidosite Diabase- Act, Chl, (Ep) Diabase- Unclass

0.704778 0.706128 0.705524 0.705415

28 30 22 26

71.0 81.1 41.6 37.9

0.1 2.6 0.1 0.1

0.704774 0.706009 0.705517 0.705402

2220 2018 1793

49140-387245 49080-387315 49075-387435

Diabase- Act 0.706139 Diabase- Act, Chl, (Ep) 0.706292 Diabase- Act, (Chl) 0.705002

58 38 28

110,6 151.2 73.7

4.3 2.3 3.7

0.705995 0.706235 0.704817

DiabaseDiabaseDiabaseDiabaseDiabase-

77 26 42 54 30

135.9 77.4 121.7 78.3 113.6

3.9 0.4 1.3 0.9 0.9

0,705695 0.705491 0.705283 0.705225 0.704851

II IT IT l! It I! !I 11 II

,, Vl

f

1! T! I!

!I t~ T! IV 11 !I II I! !I IV

Sheeted Dyke Complex MM73/1 MM75 MM76/I MM77 MM78 MM79

2061 2061 2061 2061 2061 2061

51720-386670 51720-386675 51720-386675 51720-386675 51720-386675 51720-386675

Peristerona Potamos gorge traverse PP81/1 PP85/1 PP89/1

1658 1622 1720

Marathasa Valley MA292 MA294 MA296 MA298

Kakopetria K299 K300 K301

Ayios Theodhoros to Spilia ATS302 ATS303 ATS304 ATS305 ATS305/2

1988 2122 2256 2402 2463

49350-387862 49365-387688 49400-387600 49380-387500 49370-387475

Act Chl, Act,(Ep) Act, Chl Act, Chl Act,Chl

0.705803 0.705512 0.705323 0.705268 0.704882

CONSTRAINTS FROM THE TROODOS OPHIOLITE

135

Table 2. (continued) Sample number

Depth (m)

GR

Sample description

878r/86Sr

2sigma Sr ppm

Rb 87Sr/86Sr ppm (t = 90 Ma)

Nikitari to Asniou NA307 1957

49805-387855

Diabase- Act

0.705098

24

80.7

0.9

0.705055

Xyliatos to Lagoudhera X308 1579 X309 1793 X310 1970 X311 2073

50355-387575 50355-387380 50280-387215 50165-387085

DiabaseDiabaseDiabaseDiabase-

0.705393 0.705330 0.705420 0.705320

20 62 50 42

78.3 86.7 78.7 130.7

4.8 0.3 0.2 1.8

0.705165 0.705317 0.705411 0.705270

Apliki to Palekhori A317 2335 A318 2390

50965-386545 Diabase- Act 50900-386500 Diabase- Act

0.704806 0.704551

42 20

84.6 88.7

1.0 0.4

0.704762 0.704533

38 42 24 76 16 58 18

78.2 41.0 82.9 83.8 96.7 115.2 114.2

0.6 1.4 1.7 0.7 0.8 1.0 2.1

0.705299 0.705118 0.705243 0.704145 0.705739 0.704674 0.704983

14 13 13 13 14 13 16

97,2 120,6 112.4 247.6 167.3 282.0 119.6

2.3 4.8 1.7 0.0 0.4 0.3 0.1

0.705838 0.705997 0.705239 0.705158 0.705256 0.705127 0.705670

Chl,/Act Chl, Ep, Act Act, Chl,(Ep) Act

Apliki to Makheras to Kapedhes KM319 2409 51215-386395 KM321 1963 51370-386840 KM322(2) 2 0 7 3 51525-386715 KM323 2049 51595-386690 KM324/2 1 9 3 9 51915-386820 KM325 1787 52180-386925 KM325bla 1 7 8 7 52180-386925

DiabaseDiabaseDiabaseDiabaseDiabaseDiabaseDiabase-

Phterykhoudhi Potamos A22" 2230 A24" 2230 A26" 2250 A27" 2250 A35" 2300 A36" 2260 A37" 2260

Diabase- Act, Chl Diabase- Chl, (Act) Diabase- Act Epidosite Qtz-chl-ep Epidosite Diabase- Act, (Chl)

50851-387166 50851-387166 50451-387160 50451-387160 50866-387162 50878-387177 50886-387175

Chl, Act Act, Act Chl, Act, Act.

(Act)

0.705329 0.705244 Chl, (Ep) 0.705318 0.704176 (Ep) 0.705768 Chl 0.704706 Chl, (Ep) 0.705050 0.705924 0.706144 0.705294 0.705158 0.705264 0.705131 0.705673

* Analysed following methods of Teagle et al. (1996): other 87Sr/a6Sr analyses performed on VG54E mass spectrometer. Act-actinolite; Chl-chlorite: Ep-epidote, Plag-plagioclase, Cpx-clinopyroxene: Brackets indicate mineral present only in trace amounts. Qtz-chl-ep-Diabase completely recrystallized to granoblastic quartz-chlorite epidote. Descriptions of CY4 samples after Baragar et aL (1989).

Figure 3 shows the new analyses c o m p a r e d to previously published data on an 87Sr/86Sr depth profile for the basal group and sheeted dyke complex. With the exception of the CY-4 profile, the majority of the new analyses of samples from the sheeted dykes are within the previously reported range (0.7047-0.7059; Bickle & Teagle 1992). T h e r e are no significant differences between samples from different structural grabens or dyke domains. This b r o a d scale h o m o geneity strongly supports the assumptions of Bickle & Teagle (1992) and only slightly enlarges the 87Sr/86Sr range used to c o n s t r u c t their isotopic tracer model with 95% of the analyses n o w falling between 0.7045 and 0.7059. The previously published and new Sr-isotopic analyses of the regional diabase sheeted d y k e samples are c o m p a r e d to analyses of the highly altered epidosite and quartz-chlorite rocks in

Fig. 6 a n d show no systematic difference with the m e a n of 70 chlorite, actinolite-chlorite and actinolite grade samples of 0.7052 being indistinguishable from that of 26 epidosite and quartz chlorite rocks. The only exceptions to this Srisotopic h o m o g e n e i t y are samples from the CY4 drill hole which are significantly lower on average (mean 0.7045, range 0.7037 to 0.7050). The CY-4 samples are significantly less altered than most o f the sheeted dyke samples, containing actinolite, igneous plagioclase and clinopyroxene, and, as discussed below, the a m o u n t of Sr-isotope exchange depends on the extent of mineral recrystallization. Fluid c h a n n e l l i n g a l o n g faults Fluids are c o m m o n l y channelled along faults and shear zones within crustal rocks (Sibson et

136

M. J. BICKLE E T AL.

Epidosites and quartzchlorite rocks Mean = 0.7052 + 3, 1~7 (n = 26)

b)

E !

Diabase (chlorite grade only) Mean = 0.7055 :k 4, 1~ ~ (n = 10) [

Iiiiiiii~'~]

Diabase (actinolite & chlorite) Mean = 0.7052 i 5, 1t~ iiiiiiiiiiiii[ii i iiiii[iiliiiiii (n = 40) iiiiiiiii:~iliiiiiiiiiiiiiiliiiiiiiiiiiiii i!iiiiiiiiiiiUili!iii!i'~iil]iiiiiiiiiiiiii i!!iiH!i!!iii iiiiUii!iii~iiiiiiiiiiiiiiliiiiiMiiii i i~,iii!iiii

t

I

Diabase (actinolite only) ~ Mean = 0.7050 4- 5, 1~ (n = 20) [iiiiiiU~iiiil i!i':ii!iiiiliiiiiiiii,~iiiiliii~,iiiiiiiiiiil t : iiiiili!', i:,!i] I: :~i!i!!!il !iiiiiiiiiiiil d),_ [----]----q i, i~il ?, i l ?,il |iiiiii!'~ii!l

,iiiiiii!

Diabase from Hole CY-4 Mean = 0.7045 4- 7, 1~

= 26)

iiiiiiiiiiiiiiliiiiiiiiiiiiiiil Fault rocks Sheeted ~ ~ dykes ~

~

r

o

Pillows/ u p

Gabbros

0.704

0.705

0.706 87

0.707

86

Sr/ Sr

Fig. 6. Histogram of whole rock STSr/86Sr for compilation of analyses from sheeted dykes on Troodos recalculated at 90Ma for (a) epidosite and quartz-chlorite rocks, (b) chlorite-grade diabase, (c) actinolite-chlorite grade diabase, (d) actinolite-only grade diabase, (e) samples from diabase in hole CY4 and (f) fault rocks from the indicated units. In histograms a to e open symbols are samples compiled by Bickle & Teagle (1992) and shaded symbols, this study. All samples from faults analysed in this study.

al. 1975) and similar channelling of recharge fluids by faults within the upper ocean crust has been postulated by a number of authors (e.g. Cann et al. 1985; Karson & Rona 1990). There is however, little direct evidence that faults provide conduits for a significant proportion of recharge fluids and the near-basalt 87 Sr/ 86 Sr of active black smoker vents precludes significant channelling during hydrothermal recharge or else high S7Sr/S6Sr seawater dominated fluid would

break through for only moderate fluid fluxes. An average recharge flux of 5x 1 0 6 k g m 2 would transport a modified seawater Sr-isotope signal to the discharge site if the recharge fluids were focused so that they interacted with only ~10% or less of the crust in the chlorite and higher grade parts of the upper oceanic crust. A well established association does exist between sites of fluid discharge, sulphide mineralization and fault zones from both modern and ancient environments (e.g. Constantinou & Govett 1973; Kleinrock et al. 1996). Detailed microstructural analyses of fault gouge from d6collements proximal to the dyke-pluton transition of the Troodos ophiolite has identified variable fluid compositions and temperatures (Agar & Klitgord 1995; Marquez & Agar 1995). Those studies identified zeolite veins cross-cut by higher temperature assemblages indicating prograde fluid evolution. Early zeolites are of minor abundance and evidence for extensive low temperature down-flow on these systems was not described. The over-printing of low-temperature phases (e.g. clay minerals, carbonate) by higher temperature hydrothermal minerals (epidote, actinolite, chlorite and sulphides) may obliterate petrographic evidence for initial seawater penetration to depth in the ocean crust along fault zones. However, if significant volumes of recharge fluid are channelled down fault zones then the transport of seawater strontium along the faults should generate Sr-isotopic gradients perpendicular to the margins of the fault zone. Fault zones from varied structural settings in the basal group, sheeted dyke complex and plutonic complex of the Troodos ophiolite were sampled (Fig. 6) to examine whether these structures record evidence of significant fluid down-welling. Samples have been collected from sites both proximal and distant to the graben axis and zones of hydrothermal discharge. All the faults studied were moderately to steeply dipping as were the dykes disrupted by these features. Faults in an outcrop of basal group rocks that crop out in the Peristerona Potamos gorge have been studied in detail (Fig. 5) and whole rock samples of fault breccias and mineral separates from fault gouge and cross-cutting veins have been analysed from a number of other outcrops (Table 3). The faults in the basal group outcrop (locality 3, Fig. 5) comprise a pair of narrow (10-20cm) fault breccias that offset thin lava flows and pillow lavas by at least a metre (Fig. 7). The fault zones sampled in this survey show little evidence for major channelling of recharge fluids relative to the surrounding host rocks. All

CONSTRAINTS FROM THE TROODOS OPHIOLITE

137

Table 3. Strontium isotopic analyses of fault zones and associated secondary minerals from the basal group, sheeted

dykes and plutonics of the Troodos ophiolite Sample number

Grid reference (Cyprus metric grid)

R o c k type

87Sr/86Sr

2sigma

Sr ppm

Rb ppm

878r/86Sr (90 Ma)

Samples from Peristerona Potamos; depth 1400 m (locality 3, Fig. 5;see Fig. 7) PP1/1" PP1/I*L PP4/I* PP5/I* PP6/I* PP6 / 1*L PP7/I* PP7/18*L PP7/2" PP10/I*

Pp11/1 PP12/1 PP13/I* PP14/I* PP16/I* PP16/I*L PP17/I* PP17/I*L PP17/2* PP18/I* PP19/I* PP20/1 * PP218/l* PP218/2" PP220/1 * PP222" PP228/1 * PP229/1" PP229/2"

50750-387670 " " " " " " " " " " " " , ,1 " " " " 1, 1, " " " " " " ,, i, " " " " 1, ,I " " " " " " " " , 1, 1, ,, " " " " " " " " " "

L a v a - SM leachate Lava - SM L a v a - SM L a v a - SM leachate Hyaloclastite-Chl leachate L a v a - SM L a v a - SM L a v a - SM Lava - SM L a v a - SM Lava - SM L a v a - SM leachate F a u l t breccia-Chl leachate F a u l t breccia-Chl L a v a - Ox L a v a - Ox L a v a - SM L a v a - SM Calcite Lava - SM L a v a - SM L a v a - SM L a v a - SM Calcite

0.706622 0.706574 0.705401 0.705692 0.705492 0.705261 0.707207 0.707239 0.704738 0.704561 0.704476 0.705160 0.704499 0.704588 0.706613 0.706598 0.707398 0.707424 0.707680 0.704675 0.704302 0.705106 0.705726 0.707019 0.704532 0.704408 0.704930 0.705264 0.706011

12 16 12 12 14 16 12 12 12 14 24 20 12 14 12 14 14 16 36 12 14 14 24 16 14 16 14 12 12

67.5 89.4 54.9 76.0 55.2 95.5 47.4 43.2 118.7 73.6 65.3 76.1 77.6 44.9 70.2 97.6 118.7 151.1 88.3 99.2 52.4 43.4 72.9 34.2 70.1 77.3 88.1 69.3 63.1

21.5 40.0 5.7 14.8 3.0 20.0 18.8 21.7 5.1 3.5 2.0 7.9 2.6 1.4 20.5 41.4 42.6 72.4 22.4 2.3 0.4 7.7 19.7 2.8 3.2 2.0 7.4 9.2 0.8

0.705446 0.704921 0.705020 0.704970 0.705292 0.704485 0.705735 0.705383 0.704579 0.704386 0.704364 0.705454 0.704375 0.704476 0.705531 0.705029 0.706071 0.705649 0.706739 0.704587 0.704272 0.704448 0.704725 0.706712 0.704362 0.703503 0.704620 0.704774 0.705965

16 24 40 28

31 47.3 65.0 78.7

0.0 0.2 0.0 0.6

0.705039 0.705147 0.704887 0.705225

0.1

0.705426

5 m wide fault zone - Panayia Bridge to Polystipos road; depth 2150 m GI* G2 G3 G4

50655-387245 " " " " " "

F a u l t breccia-Chl,Ep Breccia clast-Chl,Ep Diabase Diabase

0.705045 0.705165 0.704890 0.705253

Fault zone cutting diabase - Panayia Bridge to Polystipos road near Stavros; depth 2300 m. PP539

50535-387030

F a u l t breccia-Chl

0.705462

12

12.2

Fault zone from eastern margin of Mitsero Graben between Gourri and Lazania; depths 1500-1800 m L486" L504" L552" L579" L583"

51555-386720 51540-386775 51470-386810 51410-386840 51410-386840

F a u l t breccia-Chl,Act Zeolite-Laum/Still Zeolite-Cpt-Heu Zeolite - C p t - H e u F a u l t breccia-Chl,Act

0.706139 0.706213 0.707530 0.707164 0.705266

14 16 14 16 14

26.2 12.4 43.5 22.2 104.2

0.9 1.0 0.5 0.1 1.3

0.706012 0.705906 0.707491 0.707139 0.705220

Breccia from the CCSP CY-2a Drill Hole-Agrokipia; depth 980 m C Y 2 a - 6 8 1 m * * * 51205-387940

Breccia clast- Ep

0.705433

30

702

0.5

0.705431

E p i d o t e - v u g filling

0.706088

44

317.0

0.2

0.706085

Apliki Fault Zone; depth 2350 m Ap-V***

51075-386570

138

M. J. BICKLE E T AL.

Table 3. (continued)

Sample number

Grid reference (Cyprus metric grid)

Rock type

87Sr/86Sr

2sigma

Sr ppm

Rb ppm

87Sr/S6Sr (90 Ma)

Agros area - normal fault within Sheeted Dykes; depth 2600 m

LMI** LM2** LM3** LM4** LM5** LM6**

50174-386450 " " " " " " " " " "

Fault breccia-Ep,Qtz Ep-Plag-Qtz vein Diabase Gabbro, Hbl, Chl Fault margin-Ep Gabbro- Hbl, Chl

0.705594 0.705710 0.705916 0.705583 0.705492 0.703886

11 10 10 14 10 11

0.704026 0.703320 0.704952

11 8 16

Agros area - fault zone within gabbro; depth 2700 m

LM12** LM13** LM14**

50189-386425 " " " "

Fault breccia Olivine Gabbro Gabbro - Hbl

Ayios Ioannis-fault zone with sheeted dykes and gabbros - massive quartz veins and sulphide mineralisation; Depth 2650 m

LM7** LM8** LM9** LM10** LM11"*

50204-386200 " " " " " " " "

Gabbro Mylonite Ultracataclasite Gabbro- Hbl Diabase

0.704677 0.704742 0.705487 0.706167 0.705089

16 13 16 18 1l

* 87Sr/86Sr analysed on T40 VG Sector 54. ** analysed following methods of Teagle et al. (1996); *** Teagle (1993). SM-smectite-chlorite mixed layer; Chl-chlorite; Ox-Fe(O,OH)x; Ep-epidote, Act-actinolite; Laum/Stil laumontite/stilbite; Cpt-Heu-Clinoptilolite-heulandite; Qtz~luartz; Plag-plagioclase; Hbl-hornblende.

the fault gouges and breccias inspected and analysed comprised greenschist facies alteration minerals (chlorite ± actinolite + quartz + epidote + albite 4- sulphides) with low temperature minerals (calcite, mixed-layer clays, zeolites) only present in late-stage cross-cutting veinlets. Sr-isotope analyses of the gouge, breccia clasts and mineral separates (epidote) indicate that fluids associated with these structures were within the range of pervasive recharge documented by Bickle & Teagle (1992; Fig. 7) and no gradients in Sr-isotopic ratios perpendicular to faults were recognized. The exception to this are the analyses of a structure within olivine gabbros near Agros (LM 12-14), where the 87Sr/86Sr of fault rock (,-~0.7040) is only very slightly elevated above primary igneous values, and more primitive than the signature of strongly recrystallized wall rock nearby (Table 3). The hydrothermal Sr-isotopic signature and the greenschist facies assemblages present indicate that faults served predominantly as conduits for hot, evolved up-welling fluids. No significant gradients were recognized between the fault rocks and the neighbouring wall rocks, apart from minor (cm-scale, Fig. 8) zones of intense alteration adjacent to veining in fault

breccias. This suggests that the diabase matrix porosity and dyke chilled margins were sufficiently permeable for recharge flow to the base of the sheeted dyke complex to prevent fault zones dominating the recharge fluid flow geometry. It is not possible to entirely preclude the possibility that the record of early low temperature flow of seawater-like recharge fluids down the fault zones was obliterated by later upwelling hydrothermal fluids. Late-stage vein minerals have highly elevated near seawater-like 87Sr/86Sr (Table 3), 5180 values (Beynon 1996) and the occurrence of sodic zeolites (e.g. clinoptilolite-heulandite) in the basal group and sheeted dykes only within fault zones, indicates that some faults did provide access to late-stage seawater-derived fluids that were unable to percolate into unfaulted parts of the ophiolite. The extent of low temperature alteration and, by inference, fluid flux, must have been insignificant. From the intersection of the mineral Rb-Sr evolution trends with the 87Sr/86Sr seawater curve (Jones et al. 1994) this late stage fluid infiltration must have occurred at least 9-16 Ma after the formation of the Troodos ophiolite (Beynon 1996).

CONSTRAINTS FROM THE TROODOS OPHIOLITE

139

Table 4. Strontium isotopic' analyses of pillow lavas, dykes and secondary minerals from Locality 1, Akaki Canyon

(Grid Reference 51480E, 387570N); depth llOOm Sample number

87Sr/S6Sr

Rock type

2sigma

Rb

Sr (ppm)

(ppm)

S7Sr/86Sr (90 Ma)

96.6 74.4 82.6 77.7 56.1 77.6 89.0 63.9 1777.0 95.5 87.7 97.7 110.9 106.2 181.2 107.0 89.0 135.4 98.9 108.9 118.7 119.8 116.6 105.7 1222.3

3.1 7.6 2.2 2.0 2.2 2.3 2.6 1.2 19.7 3.6 1.4 2.7 4.8 5.2 17.2 36.0 27.5 5.6 34.3 0.4 10.0 3.3 14.7 20.2 3.5

0.704274 0.706589 0.704133 0.704186 0.704132 0.704286 0.704128 0.704142 0.706619 0.704089 0.704234 0.704117 0.703750 0.703985 0.703833 0.703489 0.703666 0.706001 0.703815 0.706654 0.705810 0.705984 0.704024 0.706320 0.706684

89.3 91.4 63.5 68.4 63.6 61.7 88.2

15.1 4.7 3.0 1.3 3.6 1.1 16.9

0.703814 0.704329 0.704255 0.704308 0.704011 0.704251 0.704336

Pillow lava profile - Fig. 9A

AK41/1 AK41/2 AK42/1 AK42/2 AK42/3 AK42/4 AK42/5 AK42/6 AK42/7 AK43/1 AK45/1 AK47/1 AK47/2 AK47/3 AK47/4 AK47/5 AK47/6 AK48/1 AK48/2 AK48/4 AK55/1 AK55/2 AK55/3 AK55/4 AK56/1

Lava- Sap Amyg- Mord Lava- Sap Lava- Sap Lava- Sap Lava- Sap Lava- Sap Lava- Sap Amyg- Cpt-Heu Lava- Sap Lava- Sap Lava- Sap Lava- Sap Lava- Sap Lava- Sap, Cel Lava- Cel Lava- Cel Lava- Sap Lava- Cel Mordenite Glassy margin Glassy margin Fresh glass Palagonitized g l a s s Amyg- Cpt-Heu

0.704392 0.706965 0.704230 0.704281 0.704275 0.704398 0.704237 0.704213 0.706660 0.704228 0.704294 0.704220 0.703911 0.704166 0.704184 0.704733 0.704808 0.706154 0.705099 0.706669 0.706122 0.706085 0.704491 0.707028 0.706694

36 52 22 24 24 36 20 42 26 22 30 24 174 24 18 26 20 22 20 22 24 22 16 16 22

0.704438 0.704521 0.704433 0.704377 0.704219 0.704319 0.705047

30 38 36 30 20 32 50

Dyke-Pillow profile - Fig. 9B

AK29/1 AK30/1 AK32/1 AK33/1 AK34/1 AK35/1 AK36/1

Lava- Sap Lava- Sap Dyke- Sap Dyke- Sap Dyke- Sap Dyke- Sap Lava- Cel

All 87Sr/86Sr analyses on VG54E mass spectrometer. Sap - saponite/celadonite; Cel - celadonite; Amyg - amygdale; Mord - mordenite; Cpt-Heu - ctinoptiloliteheulandite

Kinetic controls on fluid-rock

exchange

The range of S7Sr/S6Sr isotope ratios in the extrusive volcanic rocks on Troodos and the low 87Sr/86Sr ratios exhibited by the CY4 samples, which contain more of their original igneous mineralogy than typical sheeted dyke samples from Cyprus, imply Sr-isotope fluid-rock exchange was kinetically limited in the extrusive series and also in parts of the sheeted dyke complex. The nature of these kinetic controls is important for modelling geochemical transport through the ocean crust. For example, Bickle & Teagle (1992) presumed that the intermediate

and limited range of Sr-isotopic compositions of the recharge fluids in the sheeted dykes was set as the recharge fluids penetrated through the overlying extrusive series. The fluid-solid exchange kinetics may be rate-limited either by: (1) diffusive exchange between minerals and fluid; (2) the extent of mineralogical recrystallization (i.e. original minerals remain unaltered but new minerals grow in isotopic equilibrium with the pore-fluid phase); or (3) diffusion or advection of elevated 87Sr/86Sr away from fluid channels.

140

M . J . BICKLE

ET AL.

FAULT ZONE 1 "h South

~

/

Glassy

/ Megapitlow 1

IPP22, ~.r m ~

structures

/~/~

margin £ 7 ~--.. / / / ~

fl

\

~

/ ~

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~ d ~'/

¢

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-'\

,I,'5metres

/

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~ow4 .~"Z ~:V . -f' ~ •~-,,4/'0..,.,

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\

~

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~~

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F FL-~i' WI' t/~ ' "~

~

BRECCIA/

" ~

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j

' LC~W////K ~ 160V centimetr;s F-'K 13 60 centimetres~

II I

/ ~

" ~ ~~ ~ I View looking to west J

View looking to south

Fig. 7. Detailed sampling of two small faults within lavas of the Basal Group in the Peristerona Potamos gorge (locality 3, Fig. 5). Redrawn from field sketch.

0.708 87

86 Sr/

Sr

::: ~~I):: 90MaSeawater

Fault /

= : O

2one-

"................

t,q i

(90 Ma)

- ..........

2

~::.~ .i!. .......... v(~i~ v.~

0.706

x

i:!:!i71::~i~71

i',Ti',!',i'~iJ'~!',i

!

.::::::::v ................

iiiiiiiTiii:!ii! i

.

,!ii i ii "

0.704 Fresh Glass ,

I

100 • X

•

,r

~7~:ii . . . . . . . . . . . . . . . . . . . . . . . . . ............... ..... ........ ......... .....

::::::::t::i 200

,

I , I , I , 300 400 500 Distance along outcrop (cm)

Smectite-chlorite lavas Chlorite-beating fault & hyaloclastic breccias

I 600

© Amgydale calcite [] Amygdale - leached residues

Fig. 8. 87Sr/86Sr profile at 90 Ma of whole-rock samples, leached whole-rock samples, fault breccias, zeolites and calcites from amygdales and leached residues from amygdales plotted against distance perpendicular to two faults illustrated in Fig. 7.

CONSTRAINTS FROM THE TROODOS OPHIOLITE /

~

~

-/

-//

, -//

~~o ° ~ "i~ " o "-I 0

,

~3[

f

O Celadonil~matrix

~ oj" ~

/

f~} ~'~4@/

0.5metres Hyaloclastite breccia

Zeolite-filled amygdale ~ /

~ ~ /3¢""~-

o

-_%~_o_:........... Vesi~ularlayer ,~_ "

A)

~ 046 047

50

o ~ o"~" 5"2% " ~ / ' ~ / - ~ . _ _ ~ '),~'[~

o

veslcmes

Limitofvisible ~- X X thermal . . . . 1 .72] Dyke X

/,~:>

"~
~- /,'-

o

141

o o ~oo--~-~ _30 , ]

o o o

~

o.'

B)

x

" :" C ,, :--j )

/:. I

x \, veins

X,~,jt~ x

X

water level 1 metre

x

'21

"

x/ ,,.~-_,~--

t • Celadonite matrix

" ~ ~ l o c l a s f i . t e breccia ¢,i"),; , ~ between pillows

Fig. 9. Detailed sampling profiles at locality 1 (Fig. 5) in Akaki Canyon. Note vesiclerich upper zones of pillows and patchy development of celadonite. 9A shows sample profiles across pillow and 9B across pillow cut by a small dyke. We have attempted to distinguish between these three possibilities by searching for isotopic gradients away from possible fluid channels (faults, flow and pillow margins) and examining the relationship between Sr-isotopic alteration and recrystallization of the original igneous assemblages.

Extent of fluid channelling on flow and pillow margins: Akaki Canyon, Locality 1 Fluid flow might be expected to be focused in the more glassy and fractured margins and interpillow hyaloclastites of the pillow lavas and flows (Gillis & Robinson 1990) and along other potentially more permeable pathways such as the faults discussed above. Evidence for such channelling of fluid flow has been sought by looking for gradients in 87Sr/86Sr ratios away from pillow, flow and dyke margins. At Locality l, three layers of gently northward-dipping lava pillows are cross-cut by several subvertical dykes less than 1 m thick. The pillow lavas are heterogeneously altered to secondary minerals, predominantly Mg-saponite smectite with the patchy development of deep green celadonite-rich zones (Fig. 9A). Five to ten centimetre thick glassy margins separate adjacent pillows. Plagioclase mostly retains igneous compositions with anorthite contents between 65 and 79 reel. % with some less anorthite-rich grains (15-65 mol. %) and rare albite grains. Clinopyroxene phenocrysts and magnetite are little altered. Within the glassy margins of the pillow lavas palagonitic envelopes are developed along orthogonal crack networks around islands of vitreous, isotropic glass.

The cross-cutting dykes (Fig. 9B) display similar alteration assemblages to the pillow lavas. A 15cm-wide celadonitic halo is developed within the pillow lavas adjacent to the chilled margins of the dyke. Petrographic observations and oxygen-isotopic constraints (Beynon 1996) indicate that the celadonitebearing assemblage overprints the earlier development of saponite. A similar paragenesis is observed in drill-cores into old ocean crust (e.g. in ~100 Ma crust, Alt et al. 1992; Alt 1993). A Sr-isotopic profile across the pillow in Fig. 9A shows that most of the samples have 90 Ma 87Sr/SOSr ratios on the upper boundary or slightly above the range of the least altered basaltic glasses (Rautenschlein et al. 1985, Table 4; Fig. 10A). Rb/Sr ratios of pillow samples are sufficiently low that the estimated 87Sr/S6Sr ratio is insensitive to the exact timing of the alteration (Table 4). Celadonites have high Rb/Sr ratios and the significance of their 90 Ma 87Sr/S6Sr ratios are uncertain because the mineral may have crystallized significantly after formation of the crust and there is evidence of subsequent perturbation of Rb-Sr and K - A r isotope systematics (Booij et al. 1995). There is no significant 87 Sr/ 86 Sr gradient from core to rim of the pillow. The surrounding glassy margin, altered glass separates and amygdale infills yield significantly higher 87Sr/86Sr ratios. Altered glass and whole rocks have similar strontium concentrations, suggesting that the differences in S7Sr/86Sr are due to the proportion of fluid and rock Sr exchanged and not simply greater addition of seawater Sr to the altered glass. Amygdale 87Sr/86Sr ratios (0.7066-0.7070) are slightly lower than Cretaceous seawater, requiring the incorporation of minor basaltic strontium.

142

M. J. BICKLE E T AL.

Similarly, there is no obvious Sr-isotopic gradient across the dyke intruded into the pillow lavas (Fig. lOB). The alteration assemblages observed (saponite versus celadonite) and the possibility that these alteration stages occurred at different times makes assessing the scatter of the data difficult, though a slight heterogeneity in 87Sr/S6Sr is seen for samples with similar petrographic appearances and secondary mineral assemblages. The uniform Sr-isotopic composition across both the pillow and the dyke strongly suggests that transport of the Sr-isotopic signature away from obvious potential high flow channels was not the rate-limiting control on the exchange kinetics. The rather low 87Sr/S6Sr of whole-rock samples of the pillow interiors (~-,0.7035 to 0.7043), less than that of the underlying altered sheeted dyke complex, coupled with the high 90 Ma SVSr/S6Sr ratio of zeolites in amygdales and of glass, implies that the pillow interiors are only partially exchanged with the local pore fluid. It is probable that the pillow mesostasis comprises a mixture of igneous phases (clinopyroxene and Ca-plagioclase) with basaltic87Sr/8-6Sr ratios and saponite and celadonite altered to St-isotopic equilibrium with the pore fluid. The samples are too fine-grained to allow separation of clay mineral and igneous phases to test this hypothesis, but the alteration assemblages from the basal group and sheeted dyke complexes discussed below suggest that unrecrystallized igneous minerals preserve their original 87Sr/86Sr ratios. This situation is observed in modern ocean floor rocks (Teagle et al. 1996, this volume). The mineral modes in the pillow samples AK42, 43, 45 and 47 are consistent with the whole-rock 87Sr/86Sr ratios being a mixture of igneous phases with typical Sr contents (clinopyroxene 50ppm, plagioclase 100 to 200ppm) and basaltic 87Sr/86Sr ratios of ~0.7036 and celadonite ( < 2 0 p p m Sr) or smectite (10-50ppm) with S7Sr/8-6Sr between 0.705 and 0.707 (Beynon 1996). The homogeneity of the Sr-isotope profile across the pillow and dyke at Locality l may be because: (1) fluid flow was relatively homogeneous and not channelled on pillow, flow or dyke margins; (2) diffusive or advective transport of Stisotopes away from fluid channels was sufficiently rapid to homogenize the 87Sr/S6Sr of the pore fluid; or (3) diffusive exchange in the pore fluid took place over a protracted time interval after the high-flux circulation ceased.

The Rb-Sr isotope systematics of both zeolite phases from amygdales and pillow interior samples are plotted on an isochron diagram in Fig. 11. The scatter of analyses within the two sample groups (zeolites and pillow interior samples) is too great to make meaningful age estimates. The marked difference between the 87Sr/86Sr ratios of pillow interiors and the zeolite minerals precludes significant Sr-isotopic homogenization after crystallization of the alteration assemblages as a mechanism for eliminating early Sr-isotopic gradients. The lack of a significant difference in the 90 Ma 87Sr/86Sr ratios between zeolites collected from the interstitial material between pillows and from amygdales within the pillows also indicates lack of Srisotopic gradients within the pillow at the time the zeolites crystallized. The zeolites may have crystallized early during the penetration of the recharge fluids which fed the high-temperature flow or later during the waning phases of the hydrothermal circulation. K - A t and Rb-Sr ages on celadonites from the lavas on Troodos range from ~90 Ma to as young as 50 Ma (Staudigel et al. 1986; Gallahan & Duncan 1994; Booij et al. 1995) and have been cited as evidence that fluid circulation continued for ~40 Ma after formation of the crust. This is consistent with the observation that the conductive component of oceanic heat flow is low in crust up to 60Ma (Stein & Stein 1994). However, the low temperature flows which perturb heat flow may circulate only in the uppermost part of the lavas and have little or no geochemical impact on the crust. The K - A r celadonite ages are systematically younger than Rb-Sr ages on Troodos which implies that celadonite is open to exchange after crystallization, a conclusion that was confirmed by ion-exchange experiments (Booij et al. 1995). In addition, celadonite is one of the lowest temperature of the alteration minerals, forms in oxidizing conditions, and probably crystallizes very late in the alteration sequence. Thus the celadonite ages only put a maximum limit on the duration of the hydrothermal circulation on Troodos. However, since 87Sr/S6Sr ratios in the fluid at any depth are likely to increase with time as the elevated S7Sr/86Sr ratio of seawater penetrates deeper into the crust and as fluid-solid exchange rates decrease as the crust cools, the 87 Sr/ 8o~Sr ratio of the zeolites place an upper limit on the local fluid SYSr/86Sr of ~0.7067. This is close to the fluid composition predicted by the modelling of Bickle & Teagle (1992) and within the range of possible SVsr/a6sr ratios of the smectite compo-

CONSTRAINTS FROM THE TROODOS OPHIOLITE 87 0.708 i_Pi}low- t Pillow 878r/86Sr i - -~ . . . . . . . . . . . . . . . . .

(90 Ma) !

"D

90Ma seawater

o

I Saponite/celadonite [] Celadonite A Altered glass o Zeolite

, z~ ]~

0.706

Glassy_t Margin -':1-"

iiiiiiiiililii '

0

'

A)

'

~

20

,

iiiii~iiii::iiiiiiiiiiiiiiiiiiiiii

"

"

"

i

40

,

,

,

i

60

,

,

,

Distancealong outcrop (cm)

i

,

~ <

Dyke

>r='~ Pillow

-it

IiiJEvel°p2i! ~:-i

Celadonitic

"'.

,

~--~

....

Fresh il

iiii!~iiiii!iiiiiiii!!iiiiiiiiiiiig~a~s ,

Pillow Lava

0.705

0.704

1

0"7041

86

Sr/ Sr(90 Ma)

143

i! ,

80

,

100

0.703

,

,

B)

,

,

20

,

40

,

i

60

. . . . 80 100 120

Distancealong outcrop (cm)

140

Fig. 10. 87Sr/86Sr profiles recalculated to 90 Ma across the pillow and pillow/dyke outcrops sampled at locality 1 and illustrated in Fig. 9. (A) illustrates data from pillow in Fig. 9A projected onto vertical section. (B) Illustrates data from pillow and dyke illustrated in Fig. 9B projected onto horizontal section. Note that whole rock samples are only slightly elevated compared with range of fresh glass analyses (after Rautenschlein et al. 1985) and zeolite minerals have lower S7Sr/S6Sr ratios than late Cretaceous seawater (0.7074, Jones et al. 1994). 90 Ma 87Sr/S6Sr ratios of celadonite are of uncertain significance since celadonite may have crystallized significantly younger than 90 Ma and their Rb-Sr and K Ar isotope systematics have been perturbed post-crystallization (Booij et al. 1995). nent of the pillows presuming they comprise a mixture of altered clay minerals and unaltered igneous minerals as discussed above.

D y k e intruding basal group lavas A k a k i Canyon, Locality 2 At a second locality in the Akaki canyon, close to the contact between the Lower Pillow Lavas and the basal group (Locality 2; Fig. 5), a series of sub-parallel dykes (0.75-1.75m) intrude a massive lava flow (Fig. 12). Laumontite-filled vugs indicate that this outcrop is within the Transition Zone of Gillis & Robinson (1990). The heterogeneous nature of alteration is apparent in outcrop, with ~10 cm-wide pale grey, margin-parallel stripes developed within the dykes. These stripes are restricted to within 55 cm of the dyke margins, and are absent from the dyke cores and lavas. A relict igneous texture is preserved by both the dykes and lavas. Turbid plagioclase laths are partially to completely altered to albite or rarely K-feldspar. Clinopyroxene and the mesostasis are replaced by phyllosilicates. Amygdales and vugs are filled by (in the order of crystallization) phyllosilicates, laumontite :t: opaques, and calcite-quartz. Samples from the massive lava and the pale stripes within the dykes are more strongly recrystallized and comprise the assemblage albite + chlorite. Phyllosilicates from other parts of the dykes are mixed-layer chloritesmectites and plagioclase is only partially recrystallized to albite (Beynon 1996). Phyllosi-

licate composition and the homogeneity of albitization are the only mineralogical features that distinguish the pale bands from the darker portions of the dykes. A SVSr/S6Sr whole rock profile across the massive lava and two adjacent dykes (Fig. 12) shows a bimodal distribution, with samples that display complete recrystallization to albite and chlorite yielding significantly higher ratios than the partially recrystallized chlorite-smectite bearing-dykes (87 Sr/ 86 Sr ~0.7050-0.7052 versus 0.7045-0.7047, respectively). Across the outcrop there is a very slight decrease in S7Sr/S6Sr from the lavas into the chlorite-smectite rocks that comprise the core of the dykes. A late stage laumontite vug filling within the massive lava has a S7Sr/S6Sr=0.706427, significantly higher than the host lava. S7Sr/S6Sr v. 1/Sr mixing relations (Fig. 13) show that the elevated a7Sr/S6Sr of the lavas and pale stripes cannot be accounted for by contamination by late-stage laumontite or other phases. The differences in Sr-isotopic ratio between the two alteration assemblages reflect the presence of relict igneous minerals within the chlorite-smectite rocks. The similarity of the S7Sr/86Sr of 100% recrystallized lavas and stripes to that of the quartz-chlorite and epidosite rocks as well as diabase from the sheeted dykes is evidence that these albitechlorite rocks preserve the isotopic signature of the fluid responsible for the alteration. A fluid with 87Sr/S6Sr ~0.7050~).7052 is in the middle of the range of compositions suggested for recharge fluids in the sheeted dykes (Bickle & Teagle 1992).

144

M.J. BICKLE ET AL. Key

S7Sr/86Sr 0.7070

..........

Zeohtes Age = 73 + 1 6 M a IR = 0.706656 4- 34 M S W D = 1.3

~B .... X

0.7060

+

X

• • * + ×

glass & pillow margins Smectite bearing samples

0.7050

AK47 Age = 57 + 22 M a IR = 0.703999 :t= 166 M S W D = 32

÷

Celadonitebearing samples

0.7040 • 0.0

AK42 AK41 AK43 AK45 AK47 AK48 AK55 AK56

87Rb/86Sr 0.2

0.4

0'.6

().8

1'.0

Fig. 11. Zeolite and whole-rock samples from pillow locality 1 (Fig. 9A) plotted on isochron diagram. Zeolite minerals and smectite and smectite-celadonite whole rock samples have very different Sr-isotopic compositions. Zeolite samples exhibit too little dispersion in Rb/Sr ratios to define a precise age. Whole-rock samples from AK47 scatter about regression in excess of analytical error as expected for samples which are partially altered and no age significance can be attached to regression.

Table 5. Strontium isotopic analyses of massive lavas, dykes and secondary mineralsfrom Locality 2, Akaki Canyon (Grid R~:ference 51350E, 387410N); depth 1500m Sample number

Rock type

AK71/1 AKKD AK69/1 AK68/1 AK67/1 AK66/I AK66/4 AK65/1 AK63/l AK62/1 AK61/1 AK60/1 AK58/1 AK58/2 AK57/2 AK72/I AK72/2

Lava Lava- Amyg laura Lava- Sap Lava- Sap, Chl Lava- Chl Dykel- Sap Dykel- Sap Dykel- Sap Dykel- Sap Dyke 1- Sap Dykel- Sap Dykel- Sap Dyke2- Sap Dyke2- Chl Dyke2- Sap Dyke 2- Sap Dyke 2- Sap

S7Sr/86Sr

0.705244 0.706513 0.704976 0.705321 0.705299 0.704768 0.704700 0.704918 0.704799 0.704779 0.704837 0.704715 0,704548 0.705303 0.705036 0.704557 0.704179

2sigma 24 38 22 28 26 34 18 40 30 32 20 24 20 22 24 30 28

Sr (ppm)

(ppm)

87Sr/86Sr (90 Ma)

35.1 32.7 70.0 46.8 21.8 150.9 173.0 80.0 157.9 158.1 113.6 137.6 136.5 32.6 121.1 122.3 81.8

1.5 0.8 5.8 6.8 0.9 4.6 5.6 4.4 5.6 5.6 4.3 6.4 1.9 2.1 14.5 1.4 1.1

0.705083 0.706427 0.704672 0.704783 0.705151 0.704654 0.704580 0.704712 0.704669 0.704648 0.704699 0.704542 0.704495 0.705069 0.704594 0.704515 0.704667

Rb

87Sr/a6Sr analyses performed on VG54E mass spectrometer. Sap-saponite-chlorite mixed-layer; Chl-chlorite; Amyg lauln laumontite filling amygdale. The smectite-chlorite samples scatter about a 90 Ma reference line on an isochron diagram, but the more altered chlorite-albite samples lie well above this line indicating addition of a high 87Sr/86Sr component during the alteration (Fig. 14). The completely altered chlorite-albite rocks exhibit too little scatter in Rb/Sr ratios to constrain the timing of alteration. The extent of Sr-isotopic alteration at this locality is dependent on the degree of recrystal-

lization of the igneous mineral assemblage. What causes this variability is less easy to determine. It may be due to the passage of different fluid volumes through the matrix porosity of the outcrop, with greater fluxes flowing through the more permeable lavas and dyke margins resulting in more extensive recrystallization, higher 87Sr/86Sr and lower Sr concentrations (Table 5). If fluid had been channelled along igneous boundaries then an

CONSTRAINTS FROM THE TROODOS OPHIOLITE Lava

87 86 Sr/ Sr 90 M a 0.7065

Dyke

145

Dyke 2 )rite-albite (An
0.7060

ctite/chlorite Plagioclase 5 to 63%)

0.7055

ed chlorite/ ctite/chlorite mblage 1 to 8%) montite 1amygdale

0.7050

0.7045 -200

-100

0

100 200 300 Distance cm Fig. 12. 87Sr/S6Sr profile recalculated to 90 Ma across dykes intruding lava flow close to contact between basal group and lower pillow tavas (locality 2, Fig. 5). Shaded zones in dykes indicate presence of ~10 cm thick pale gray margin parallel zones of more altered albite~hlorite assemblages.

isotopic gradient would have developed perpendicular to the conduit. However, there is only a very slight 87Sr/86Sr gradient across the two dykes and assemblages principally vary parallel to the dyke margins. It may be that the chloritealbite forming reaction and associated metasomatism involves a volume loss which causes matrix fluid-flow to become organized in highflux layers whose spacing is a complex and unknown function of rock reaction and compaction rates. The implication is that matrix porosity flow dominated alteration at this outcrop with the rate of secondary mineral formation, as indicated by the completeness of albitization and chlorite crystallinity, controlling the extent of fluid-rock isotopic exchange. Controls on Alteration in the Sheeted D y k e Complex Although most of the sheeted dyke complex is altered to relatively homogeneous S7Sr/86Sr ratios between 0.7047 and 0.7060 (mean 0.7052 + 5, lcr; Fig. 6), the less altered actinolite, clinopyroxene and igneous plagioclase-bearing dykes from drill hole CY4 have lower 87Sr/S6Sr r a t i o s b e t w e e n 0.7035 to 0.7050 ( m e a n 0.7045 + 7, lcr). The main mineral reservoir for Sr in the holocrystalline dyke samples is plagioclase and the relationship between recryst a l l i z a t i o n of plagioclase and w h o l e - r o c k 87Sr/S6Sr ratio for regional samples of the sheeted dykes is illustrated in Fig. 15 where the

mean and range of plagioclase anorthite contents are plotted against whole-rock S7Sr/S6Sr ratio. However, the good correlation between mean plagioclase An content and whole-rock 87Sr/86Sr ratio implies that the plagioclase grains within individual samples also probably exhibit a range of 87Sr/86Sr ratios which correlate with their An contents. An implication might be that plagioclase recrystallized at progressively lower temperatures to lower anorthite contents and higher S7Sr/86Sr Sr ratios which reflects cooling of the rock and increasing 87Sr/86Sr in the recharge fluid. Since St-isotopic exchange with the fluid is controlled by the amount of recrystallization, the factors which might control this are of interest. The degree of recrystallization is not controlled by temperature as all the rocks cool from igneous temperatures; nor, given the degree of Sr-isotopic alteration, is simple availability of water to form hydrous phases a likely control. More probable controls include: (1) variations in the local fluid flux given that areas of high flux would have higher porosity allowing fluid more access to grain boundaries and facilitating diffusive exchange between reactant and product minerals; or (2) the previous alteration history as regions initially altered at low grade might be more prone to complete recrystallization if subsequently heated in the presence of a substantial fluid flux.

146

M.J. BICKLE E T AL.

0.7065

"""

87Sr/86Sr 90 Ma 0.7060

" .-'"

, Void-filling secondary minerals

" ~ / ...... _~ 'v~' ~' /

Chlorite, complete albitisation ... / ............. -~-............ ,/"" [] .-" "" ~. [] [] .............................

. . ~,,~)~ Smectlte/chlorlte ~ incomplete ' ~'/ albitisation .W~-~/ , @,~

0.7055

0.7050

~. . . . , ,'il/ii 8I • ) /, E ~,~m • ,,

0.7045

,'---

0.00

.+'/

[] AK68 partially chloritised 1

L

0.01

I

L

0.02

l/Sr

I

,

0.03

,

0.04

0.05

Fig. 13. S7Sr/S6Srat 90 Ma versus 1/Sr for samples from locality 2. Chlorite-albite samples do not lie on a mixing line between saponite-bearing samples and compositions of void-filling minerals indicating that alteration involves exchange of Sr between fluid and recrystallizing minerals and not simply addition of new alteration phases. 0.7054 87

86

Sr/ Sr /'[3

D'"

•

0.7052 Chlorite-albite samples

c

~

/

0.7050 /z / . ~ /

0.7048

]

/

• /

Smectite/chlorite, partially albitised samples

•

0.7046 Eli ,

0.0

I

0.1

,

I

,

0.2 87

I

86 0.3

,

I

0.4

Rb/ Sr

Fig. 14. Isochron diagram of samples from locality 2. Although saponite-chlorite samples scatter about the 90 Ma reference line, more completely altered chlorite-albite samples lie away from this line indicating addition of a high 87Sr/86Sr component.

The second possibility seems unlikely as there is no evidence that the dyke complex underwent initial recrystallization at low grades (dyke samples typically preserve igneous textures and there is evidence of chlorite replacing actinolite in samples transitional to greenschist facies). Also reheating of the whole crust would involve intrusion of a layer several kilometres thick within the underlying cumulate rocks.

Reconciling the thermal and Sr-isotopic profiles The widespread c h l o r i t e - albite -4-epidote or actinolite-chlorite-Ca plagioclase assemblages in the sheeted dyke complex imply recrystallization temperatures in excess of ~200°C and probably 250°C (Alt et aL 1996a; Laverne et

CONSTRAINTS FROM THE TROODOS OPHIOLITE al. 1995; Vanko et al. 1996). The Sr-isotopic alteration over the ~1 km thick sheeted dyke zone implies that the time integrated fluid flux was sufficient to transport the altered Sr-isotopic signal more than 1 km during this alteration (Bickle & Teagle 1992). The Sr-isotopic composition cannot have been set by subsequent lower temperature alteration because:

147

through the ~2 to 2.5km thick hydrothermal system is supplied from the underlying ~5 km of crust and the hot hydrothermal fluid exits up narrow pipe-like structures which lose limited amounts of heat to the surrounding crust. Such an approximation compares well with twodimensional simulations (e.g. Sleep 1991). In the models discussed below, the temperature at the base of the hydrothermal system during the recharge phase is presumed to be 400°C, that of the source of many black smoker fluids. In such a one dimensional hydrothermal system, the initial thermal profile (here taken as linear gradient from zero at the ocean floor to 400°C at its base) is cooled by the downwards flow of the recharge seawater. The thermal structure evolves to an equilibrium state with the upper part of the crust close to the temperature of the infiltrating seawater and higher temperatures confined to a boundary layer at the base of the system (Fig. 16 a & d). In the boundary layer, heat supplied from the underlying heat source (a convecting magma chamber at ocean ridges) compensates for that lost by advection. The thermal structure can be described by one dimensionless number, the Peclet number (Table 1), a function of the fluid flux, the length scale of the system and thermal diffusivities of fluid and rock (see Bickle & McKenzie 1987). 7 The time-integrated flux of 3 x 10 kg m -2, inferred from the Sr-isotope profile of the Troodos ophiolite, implies alteration with a Peclet number of ~12. This calculation assumes

(1) zeolite facies or other sub-chlorite grade assemblages are rare and generally have seawater-like 87Sr/86Sr ratios; and (2) actinolite and epidote would not exchange Sr with a fluid phase below their blocking temperatures (> 500°C for actinolite, Cliff 1993) and these phases have similar S7Sr/S6Sr ratios to the recharge fluids as shown by the Sr-isotopic compositions of the epidosites. The sheeted dyke complex must therefore have been maintained at temperatures in excess of o . . . . . 00 C while subjected to a tlme-mtegrated firedflux of > 3 x 107 kg m -2, that is required to alter the Sr-isotopic composition. This observation is inconsistent with simple thermal models for evolution of the oceanic crust. The thermal structure of the oceanic crust is time-varying and three-dimensional. However, a one-dimensional model, with steady state or episodic recharge flowing vertically downwards over a 2.5 km section, illustrates the important features. This is because most of the heat lost 0.706 Minimum range of fluids in -Dykes 0.705 87

-4-

t O

86

Sr/ Sr 90 Ma 0.704

iiiiii!iii!iiii!ili!iliiiii

0.703

,,, 0

I,,, 20

Fresh Glass ~ii-ii-!~-i~i/ I , , , I , , , I ~ , , I 40 60 80 100

Anorthite % Fig. 15. Whole-rock 878r/86Sr ratio, recalculated at 90 Ma, plotted against mean and range of plagioclase anorthite contents for samples from sheeted dykes. Diagram implies a correlation between plagioclase An mole % and 878r/86Sr and implies that plagioclase grains within samples may exhibit a range of 87Sr/S6Sr ratios with a similar correlation between mole % An and 87Sr/86Sr ratio. Data from Beynon (1996), Teagle (1993) and glass data from Rautenschlein (1987).

148

M.J. BICKLE E T AL.

that the hydrothermal circulation persisted for 200ka, to simulate recharge extending 2kin from the ridge in crust spreading with a 1.25 cm a -a half-spreading rate. In a system with a Peclet number of 12, the basal boundary layer in which temperatures would be in excess of ~200°C would be only ~ l S 0 m thick (Fig. 16a). Hydrothermal circulation at such a slow spreading ridge would probably be intermittent (c.f. the geochronology of TAG; Lalou et al. 1993), but intermittent circulation with transient high fluxes gives essentially the same thermal structure, albeit with some temperature increase related to conductive heating during times of no flow (Fig. 16b,c). It is clear that normal thickness oceanic crust cannot supply enough heat to drive sufficient fluid flow to explain the extent of St-isotope alteration at temperatures above 200°C as discussed above. It is interesting that more modest fluxes of ~5×106 k g m -~, (equivalent to a global rate of 2×1013kga ~), allow the basal ~1 km of the hydrothermal system to remain at temperatures above 200°C (Figs 16d, e & f). The available, albeit limited, evidence from existing oceanic crust suggests that such fluid fluxes are consistent with both the alteration temperatures and St-isotope structure of the crust. The less altered Sr-isotope profile of DSDP-ODP Hole 504B could be formed by a time-integrated recharge flux of ~5 × 106 kg m (Bickle & Teagle 1992; A l t e t al. 1996b; Teagle et al. this volume) and such a flux would vent black smoker fluid with an 87Sr/86Sr ratio close to that of basalt as observed at most active oceanic vent fields (Palmer 1992). It would be possible to alter a larger volume of rock at high temperature if the high temperature hydrothermal boundary migrates through the rock as the system evolves. However, the only circumstance in which the boundary layer would migrate slowly enough to allow the observed St-isotopic alteration would be where it overlies the top of a magma chamber which sinks down through cumulate rocks with progressive crystallization on its roof. But this mechanism cannot explain the greenschist or amphibolite facies alteration of ~ l k m of sheeted dykes, which, from their grain size must have been intruded into cooled crust; nor is there evidence for significant Sr-isotopic alteration of more than a few hundred metres of the gabbro sequence in the Troodos ophiolite. There are two explanations for the alteration pattern in Troodos ophiolite: Either the alteration state of the ~ 4 0 k m long crustal section sampled is atypical of the average crust or hydrothermal circulation in the Troodos ophiolite reflects processes not operative in normal

ocean ridge settings, such as significant off-axis magmatic underplating of the upper crust.

Conclusions Further regional sampling of the sheeted dyke complex has confirmed that the majority of the samples are altered to 90 Ma 87Sr/86Sr ratios between 0.7045 to 0.7060 (mean 0.7052 + 5, 1 a), values indistinguishable from those of the highly altered epidosite and quartz-chlorite rocks (mean 87Sr/86Sr = 0.7052 4- 3, 1 o-). This implies; (1) that much of the sheeted dyke complex was altered to Sr-isotopic equilibrium with the recharge fluid; (2) that black smokers on Troodos had elevated 87Sr/86Sr ratios (0.7052) with respect to the basaltic crust. The latter contrasts with black smokers sampled in modern oceans; and (3) that the recharge to the high temperature hydrothermal system on Troodos had a time-integrated fluid flux in excess of 3 x 107 kg m -2. Less-altered samples of sheeted dykes were only found in the CY4 drill core and a few surface outcrops in the region nearby. These samples are atypical in that they preserve unaltered magmatic clinopyroxene and anorthite-rich plagioclase, but this does suggest that there was some spatial variability in the recharge regime. The Sr-isotopic composition of the recharge fluid is set to ~0.7052 during passage through the low and intermediate temperature pillow lavas and transition zones in which fluid-solid Sr-isotope exchange is kinetically limited. Analyses across pillows, dykes and lava flows show that the alteration is remarkably homogeneous and that, if fluid was channelled to any significant extent, advective or diffusive transport away from channels was sufficiently rapid to homogenize Sr-isotopic compositions or the channels were close-spaced. The marked difference between the Sr-isotope compositions of zeolite phases and samples from pillows, lava flows or dykes is consistent with Sr-isotopic disequilibrium between whole-rock samples and fluid and precludes significant post-hydrothermal homogenization of 87 Sr/ 86 Sr ratios. It is not possible to preclude the possibility that zeolites crystallized relatively late in the evolution of the high-temperature hydrothermal systems and their compositions therefore only put upper limits on fluid 87Sr/86Sr ratios.

CONSTRAINTS FROM THE TROODOS OPHIOLITE 7 -2 Uniform Flow2003Xl0ka K g . m / ~

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One-dimensional models for thermal evolution of 2.5 km crust with upper surface maintained at 0°C and lower boundary at 400°C, initial linear thermal gradient and thermal diffusivity 10-6 m2s -~. (a, b & c) for timeintegrated fluxes totaling 3×107 kgm -2 (equivalent to a global rate of 1014 kga -1) and (d, e & f) for timeintegrated fluxes totaling 5 × 106 kgm -2 (equivalent to a global rate of 2/1013 kg a-l). (a) assumes a uniform flux over 200 ka and shows the crust reaching thermal equilibrium within about 20 ka, (b) shows cooling during the first 1 ka of intermittent flow (3 × 10 kg m -~ over 20 ka). (c) shows extent of conductive thermal re-equlhbrlatlon for the 18 ka of no flow starting with 2 ka profile from (b) and assuming top 500 m of underlying crust cooled to 700°C and crust below that at 1000°C. (d) assumes a uniform flux of 5x 106 kgm -2 over 200ka and shows the crust reaching thermal equilibrium in about 40ka, (e) shows 2 ka of intermittent flow (5× 106 kg m -2 over 20 ka) and (f) shows conductive recovery from end of 2 ka interval in (e) for the same deep crust thermal structure as for (c). Fig. 16.

•

•

7

3

.

.

.

.

.

150

M. J. BICKLE E T AL.

The c o r r e s p o n d e n c e between whole-rock 87Sr/S6Sr ratio and plagioclase anorthite content in samples of sheeted dykes, w h o l e - r o c k 87Sr/86Sr and extent of recrystallization to chlorite-albite greenschist facies assemblages in samples from the transition zone, and massbalance of Sr in smectite and igneous plagioclase and clinopyroxene from pillow lavas, shows that the rate of recrystallization of the igneous mineral assemblage is the kinetic control on fluid-solid Sr-isotopic exchange. The factors controlling the recrystallization rate are poorly constrained but probably relate to the local fluid flux (and hence the local permeability/porosity structure) as neither the temperature nor the physical availability of water can be important. The most problematic aspect of alteration on Troodos is how relatively high temperatures ( > 200°C) are maintained in the ~1 km thick sheeted dyke complex while subject to a timeintegrated recharge flux in excess of 3 x 107 kg m -2, sufficient to alter the Sr-isotopic composition. Sheeted dykes in normal oceanic crust such as that drilled at DSDP/ODP Hole 504B alter at similar temperatures but are neither so altered or enriched in S7Sr/SGSr ratio. There may be some special factors which enhance alteration in marginal basins above subduction zones such as intrusion or underplating of sufficient volumes of magmas off-axis to drive continued hydrothermal circulation in the back arc crust. If so marginal basins may make enhanced cont r i b u t i o n s to o c e a n g e o c h e m i c a l b u d g e t s although it still seems unlikely that ~ 2 0 % of the ocean floor (Sclater et aI. 1980) can contribute a sufficient flux to reconcile the factor of three or more discrepancy between hightemperature hydrothermal flux estimates based on thermal models and those based on ocean geochemical budgets for St. This work was supported by NERC by grant GR9/748 and a studentship to Beynon. Invaluable assistance was provided by the Cyprus Geological Survey, especially C. Xenophontos and the Director, G. Constantinou. We acknowledge discussion of these topics with J. Cann, numerous others interested in oceanic hydrothermal circulation as well as two anonymous reviewers.

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Rare earth element mobility in a mineralized alteration pipe within the Troodos ophiolite, Cyprus D. M. WELLS, R. A. MILLS & S. R O B E R T S S c h o o l o f O c e a n a n d E a r t h Science, S o u t h a m p t o n O c e a n o g r a p h y Centre, E u r o p e a n W a y , Southampton S014 3ZH,

UK

Abstract: Rare earth element (REE) mobility has been investigated in stockwork mineralized lavas and dykes that lie in a zone of former black smoker fluid upflow within Troodos oceanic lithosphere. Most lavas from the Pitharokhoma alteration pipe have REE compositions similar to Troodos volcanic glasses, indicating that on- and off-axis hydrothermal alteration processes have not induced significant net REE mobilization. A degree of REE mobility is inferred for some lavas which are depleted in the light REE + Eu. Within the lavas, high- and low-temperature hydrothermal alteration phases analysed by laser ablation inductively coupled plasma-mass spectrometry (LA ICP-MS) display contrasting REE patterns (variably light REE enriched and depleted respectively). Thus during high- and low-temperature hydrothermal alteration of the lavas there may have been some relative loss or gain of light and heavy REEs, that did not necessarily cause any substantial net REE mobilization. Low-temperature hydrothermal phases have been identified as the major repository for the REEs in many of the altered lavas. These data suggest that much of the REE signature of the alteration pipe is not a relict of on-axis greenschist facies alteration, but a lower-temperature post-mineralisation overprint acquired in the waning stages of hydrothermal activity and during subsequent ageing of the oceanic basement. Results from investigations of modern seafloor hydrothermal vent sites, typically at constructive plate margins, have made a significant impact on our understanding of volcanic-hosted massive sulphide deposits that are widely distributed throughout the geological record. An important part of the investigation into modern vent sites has involved the use of the rare earth elements (REEs) to constrain the geochemical evolution of the hydrothermal system (Michard et al. 1983; Michard & Albar6de 1986; Campbell et al. 1988; Michard 1989; German et al. 1990; Gillis et al. 1990; Klinkhammer et al. 1994; Mitra et al. 1994; Mills & Elderfield 1995; James & Elderfield 1996). The REEs typically exist as trivalent cations of a similar size in natural systems. While this affords the group chemical coherence, redox conditions may promote the formation of Ce 4 + and Eu 2 + which can then fractionate from the trivalent REEs. Additionally, the systematic decrease in ionic radius across the group can lead to enrichments or depletions of the light REEs (LREEs: La to Eu) relative to the heavy REEs (HREEs: Gd to Lu) due to predictable differences in the way the REEs are transported in solution and partition into mineral phases. Studies of modern hydrothermal systems have elucidated numerous aspects of the ore-forming processes operating at oceanic spreading centres. Axial hydrothermal alteration of the oceanic

crust gives rise to ~350°C black smoker fluids with a REE composition that exhibits little variability between vent fields, and is distinct from either basalt or seawater (Fig. 1; Michard et al. 1983; Campbell et al. 1988; Klinkhammer et al. 1994; Mitra et al. 1994). Seawater is chemically modified during low-temperature reactions with basalt within zones of hydrothermal recharge (Hellman & Henderson 1977; Humphris et al. 1978; Ludden & Thompson 1978, 1979; Juteau et al. 1979; Alt et al. 1986; Staudigal & H a r t 1983; Gillis & Robinson 1990a,b; Minai et al. 1990; Gillis et al. 1992). However, experimental and theoretical studies have shown that black smoker fluids acquire their chemical signature at depths of 2 to 3 km within the crust, where plagioclase alteration in a reaction zone above an axial heat source generates fluids in equilibrium with greenschist facies mineral assemblages (Bischoff & Dickson 1975; Bowers et al. 1985, 1988; Berndt et al. 1988, 1989). The striking similarities in the chondrite normalized REE patterns of black smoker fluids sampled from contrasting oceanic settings indicate that reactions between heated seawater and basalt exert a primary control on the REE composition of vent fluids (Klinkhammer et al. 1994), while complexation in solution may play a significant role in subsequent R E E fractiona-

WELLS,D. M., MILLS,R. A. & ROBERTS,S. 1998. Rare earth element mobility in a mineralized alteration pipe within the Troodos ophiolite, Cyprus. In: MILLS, R. A. & HARRrSON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 153-176

153

154

D.M. WELLS

I00 ~ 10 i i i : ~ ' : ' : ' . " ' " ' "

ET AL.

ii ili : : i ~ i i i ( i ? i i ". . .

.

ii :.p~i:il ~ili lie..,ili ii ~ . .

'i

0.1 "~

0.01 0.001

"~

0.0001

~

o.ooool

0.000001 0.0000001 La

Ce

Pr

Nd

Sm Eu Gd Tb

Dy Ho

Er Tm Yb

Lu

A Average Black Smoker Fluid * Pacific Seawater "- North Atlantic Seawater -o-- N-MORB

Fig. 1. Chondrite normalized REE data for N-MORB, North Atlantic and Pacific seawater and an average black smoker fluid. Data are from Sun & McDonough (1989), Mitra et al. (1994) and Bau et al. (1996). The shaded area shows the range in REE content of Troodos volcanic glasses. Data are from Rautenschlein et al. (1985). tion (Bau 1991; Haas e t al. 1995). The application of REEs to investigations of black smoker fluid-basalt interactions at mid-ocean ridges (Gillis e t al. 1990; Bach e t al. 1996) has been limited by infrequent sampling of the lithologies in the interiors of these systems. Consequently ophiolite-based studies continue to provide important insights into those processes occurring in the sub-surface of active vent systems. The record of high-temperature alteration minerals preserved in ophiolites attests to reactions between wall rock and black smoker fluids ascending from the reaction zone, that will modify their REE composition prior to venting (Regba e t al. 1991; Gillis e t al. 1992; Valsami & Cann 1992). However, patterns of multistage hydrothermal alteration in ophiolites are the product of prolonged (~20Ma) ageing of the oceanic basement in addition to axial hydrothermal circulation at ancient ocean ridges (Gillis & Robinson 1988, 1990a,b; Staudigel & Gillis 1990; Bednarz & Schmincke 1990). Thus, in order to document the evolution of hydrothermal fluids within the oceanic crust, it is necessary to identify and geochemically characterize secondary mineral phases that have precipitated within these contrasting alteration regimes.

This study investigates REE mobility in the host volcanic rocks of a stockwork mineralized alteration pipe within the Troodos ophiolite. The distribution of REEs between high- and low-temperature secondary minerals within the lavas has been used to infer fluid compositions and mixing in a hydrothermal upflow zone, and to establish relative REE mobility during axial hydrothermal alteration and subsequent ageing of the oceanic basement.

Troodos ophiolite, Cyprus Since its recognition as a virtually undeformed fragment of Cretaceous oceanic lithosphere (Gass 1968), the Troodos ophiolite complex has been the focus of many important investigations over the last 30years. The ophiolite occupies an area of 3000kin 2 in the southcentral region of Cyprus. Differential uplift and subsequent erosion of the massif has created an annular outcrop pattern where the stratigraphically deepest rocks of the succession cropout centrally, surrounded by successively higher units (Fig. 2). A core of tectonized harzburgite and serpentinite is succeeded by cumulate peridotites and gabbros, high-level plagiogranites and massive unlayered gabbros, and an

REE MOBILITY IN A MINERALIZED ALTERATION PIPE

155

Key Ultrarnafic suite

© Majorsulphide deposit

Gabbros

• Pitharokhomadeposit

I--

Sheeted dyke complex

~

Extrusive sequence

,,,,,~ Sedimentarycover Fig. 2. Simplified geological map of the northern flank of the Troodos ophiolite complex, showing the location of the Pitharokhoma deposit and other major sulphide deposits. Inset shows location of main figure within the island of Cyprus.

extensive sheeted dyke complex. Lavas exposed at the periphery of the ophiolite complex were erupted in a supra-subduction environment (Robinson et al. 1983; Rautenschlein et al. 1985) and are overlain by ferromanganoan sediments and Cretaceous pelagic sediments (Gass 1980). Studies of the ophiolite have resulted in models for the generation and tectonic evolution of oceanic lithosphere at oceanic spreading centres (Gass & Smewing 1973; Schmincke et al. 1983; Varga & Moores 1985), and the incidence and geometry of axial hydrothermal systems (Spoon e r & Bray 1977; Spooner et al. 1977; Schiffman et al. 1987; Schiffman & Smith 1988) associated with the mobilization (Richardson et al. 1987; Schiffman et al. 1990) and concentration of metals within sulphide ore bodies (Constantinou & Govett 1973; Constantinou 1980; Adamides 1990). Further studies have emphasized the ubiquitous low temperature alteration of the upper oceanic crust that occurs during waning hydrothermalism and crustal ageing (Staudigel et al.

1986; Staudigel & Gillis 1990; Gillis & Robinson 1988, 1990a,b; Bednarz & Schmincke 1990). Troodos volcanic-hosted sulphide ore bodies vary in size from ~50 000 tonnes to over 20 Mt, and are among the closest ancient analogues of seafloor sulphide deposits. Many of them are situated on hydrothermally mineralized lavas and are demonstrably exhalative (Constantinou 1980). Stockwork mineralized lavas occur in concentrically zoned alteration pipes beneath these ore bodies and represent the channels of upwelling black smoker fluids within the oceanic crust (Constantinou & Govett 1973; Constantinou 1980). Two types of Troodos alteration pipe have been distinguished, with minor variations in secondary mineralogy attributed to lowtemperature overprinting effects (Richards 1987; Richards et al. 1989). Much of the variability in the alteration characteristics of Troodos stockwork deposits (Lydon & Galley 1986; C a n n e t al. 1987; Richards et al. 1989) is comparable to that observed in upflow-zone

156

D.M. WELLS E T AL.

Table 1. Mineralogical and geochemical characteristics of lavas within the Pitharokhoma alteration pipe* Alteration facies

Layer silicates

Other minerals

Geochemical trends

Smectitic facies

Smectitet

Fe-hydroxides, quartz, zeolites, Ca plagioclase 5, pyroxene 5, Fe-Ti oxides5

+ K, Na, Rb, Ba* - C a , Si*

Chlorite-smectite mixed layer facies

Chlorite-smectite mixed la)~er minerals~

Albite, quartz, sphene, pyrite, Ca plagioclase ~, pyroxene~, Fe-Ti oxides~

+ Na, Mg - C a , K, Rb, Ba

Chlorite-albite facies

Chloritet

Albite, quartz *, sphene t, pyrite, Ca plagioclase ~, Fe-Ti oxides~

+ Na, Mg - C a , K, Rb, Ba

Chlorite-illite facies

Chlorite*, illitet

Quartz, pyritet, anatase t

+ Mg, K, S, Rb, Ba - C a , Na, Sr

Leached facies

Illitet

Quartz t, pyritet, anataset

+ K, Ba, Rb, Fe, S - M g , Ca, Na, Sr

*Data summarised from Richards et al. (1989). tPhase always present in a given facies (although chlorite may be absent in the pillow margins of chlorite-illite facies lavas). Data from Gillis & Robinson (1988). Relict primary phase that may be present.

t

breccias recovered from the Mid-Atlantic Ridge and East Pacific Rise (Delaney et al. 1987; Saccocia & Gillis 1995), and is interpreted to reflect the H2S content of the ascending hydrothermal fluid, consistent with the measured variability of vent fluids (e.g. Von Damm 1990). The Pitharokhoma

alteration pipe

The Pitharokhoma deposit is located on the northern flank of the Troodos ophiolite close to the transition between the sheeted dyke complex and overlying lavas (Fig. 2). It is one of many stockwork-type deposits within the ophiolite, some of which are structurally contiguous with exhalative ore bodies formed on the Troodos seafloor (Constantinou & Govett 1973; Constantinou 1980). The deposit comprises 2.3 Mt of non-cupriferous disseminated pyrite and minor occurrences of massive pyrite, located within two originally vertical pipe-like zones (each ~ 1 0 0 m in diameter) of intense basalt metasomatism (Richards et al. 1989). They are comparable in lateral extent with the intensely altered and mineralized upflow zone ( ~ 8 0 m in diameter) that underlies the ~ 2 . 7 M t hydrothermally active sulphide mound in the TAG vent field located at 26°N on the Mid-Atlantic Ridge, that was recently drilled by Leg 158 of the Ocean Drilling Program (Humphris et al. 1995).

Drilling records indicate that the deposit is overlain by less altered and mineralized lavas (Richards et al. 1989). It is inferred that the deposit formed in a zone of mixing between ascending black smoker fluids and relatively unreacted seawater below the contemporaneous seafloor (Jensenius 1984; Richards 1987; Richards et al. 1989), as has been proposed for some active vent systems (Corliss et al. 1979; Edmond et al. 1979) and the mineralized 'stockwork' zone in the Ocean Drilling Program Hole 504B (Honnorez et al. 1985). This study builds on previous work which characterized the major and trace element geochemistry and alteration assemblages of the eastern alteration pipe of the Pitharokhoma deposit (Richards 1987; Richards et al. 1989). This alteration pipe is essentially intact, and one of the most completely exposed within the ophiolite. It comprises a sequence of concentrically zoned alteration facies, each defined by a characteristic secondary mineral assemblage with the most intense alteration and mineralization occurring in the pipe centre (Table 1, Fig. 3). Mineralization occurred chiefly by the preferential replacement of interstitial ferromanganoan oxide sediments, with some replacement of lavas in the centre of the pipe (Richards 1987). During hydrothermal alteration of the lavas, interstitial sediments in the peripheral

REE MOBILITY IN A MINERALIZED ALTERATION PIPE

--v------_-_ .

.

.

':"'i:::::: _ i

.

.

.

157

i~iI

; N

~•i

•

A

'i/j

--.

50 rn

Key ~

Leachedfacies Chlorite-illite facies

~

Chlorite-albite facies "--] Chiorite-smectite mixed layer facies Smeetitic facies

.....

Fault

Fig. 3. Map of the Pitharokhoma open pit with sample localities. Simplified alteration facies are fi'om Richards et aL (1989).

alteration facies were converted to hematitic jasper (Richards et al. 1989), that is found associated with all Cyprus ore deposits (Richards & Boyle 1986). Sulphide mineralogies and textures similar to modern vent sulphides have been described in sulphide scree from the western pipe of the deposit (Jensenius & Oudin 1983). Wall rock alteration

Lavas at the periphery of the pipe are altered to a smectitic assemblage (Richards et al. 1989), which outside localized zones of upwelling is typical of the pervasive low-temperature alteration of the extrusive sequence (Gillis & Robinson 1988, 1990a, b; Bednarz & Schmincke 1990) and is similar to that reported for in situ upper

oceanic basement (Alt & Emmermann 1985; Alt et al. 1985, 1986). Lavas altered to chlorite-

smectite mixed layer, chlorite-albite and chlorite-illite facies assemblages surround Mg-depleted, illitised leached facies lavas at the pipe centre (Table 1). The leached facies lavas are inferred to have reacted with Mg-deficient upwelling hydrothermal fluids (Richards 1987; Richards et al. 1989) by analogy with experimental studies (Bischoff & Dickson 1975) and modern black smoker fluids (Edmond et al. 1979). Metasomatic Mg enrichments in the chloritised lavas surrounding the pipe centre are attributed to mixing between black smoker fluids and relatively unreacted seawater at the margins of the zone of axial upflow (Richards 1987; Richards et al. 1989). The concentrically zoned alteration facies are

Smectitic

Chl-smec mixed layer Chl-smec mixed layer

Chl-smec mixed layer Chl-smec mixed layer Chl-smec mixed layer Leached Leached Leached

Lavas: 146

145 147

153 155 157 143 150 114

Hematised jasper with Hematised jasper with Granular qtz + py + Granular qtz + py +

Interstitial sediments. 111 Chl-smec mixed layer 144 Chl-smec mixed layer 149 Leached 152 Leached

void filling smec abundant qtz and colloform textured hem; occasional basalt clasts partially altered to chl-smec minor ill/ser minor ill/ser

* Alteration facies of Richards et al. (1989). fsp = feldspar; cpx = clinopyroxene; smec = smectite; mord = mordenite; cowl = cowlesite; hem = hematite; chl-smec = chlorite-smectite; chl = chlorite; cham = chamosite; qtz = quartz; ill = illite; ser = sericite; py = pyrite; sph = sphalerite.

Massive py

Twinned fsp; fracture filling zeolites (mord + cowl); fibrous smec intergrown with zeolites; occasional cracked py; abundant matrix filling smec + amorphous Fe-oxide Fe-oxide breccia with lava fragments partially replaced by equigranular qtz + chl-smec; sparse py Enlarged vesicles lined with qtz + chl-smec and filled with finer chl-smec+qtz; occasional twinned fsp partially replaced by chl-smec; amorphous vesicle and matrix-filling Fe-oxide; sparse py Fsp partially replaced by chl-smec; abundant vein and matrix-filling amorphous Fe-oxides; relict igneous texture preserved; py absent Partially altered twinned fsp and cpx; radiating orange/brown matrix and vesicle filling smec; sparse py Amorphous Fe-oxide±chl-smec±chl (cham) replacing plag and mesostasis; eroded py with minor sph; qtz + py in vesicles Relict igneous texture preserved; fsp replaced by ill + qtz; qtz in vesicles and veinlets replaces and is overgrown by py Abundant qtz + py + fine grained ill; relict igneous texture absent; vesicles contain qtz+py+ill Rare relict igneous texture; abundant py in mesostasis and associated with qtz in veinlets and vesicles; abundant fine grained matrix and vesicle filling ill

Qtz developed in mesostasis; qtz + chl-smec+chl (cham) in vesicles; Fe-oxide after py Fsp replaced by chl-smec+chl (cham); sparse py; amorphous Fe-oxide in groundmass Fsp replaced by chl-smec; qtz filled vesicles; disseminated euhedral py with minor sph; igneous texture absent in patches Albitised fsp; mesostasis replaced by chl-smec±minor chl (cham); sparse py; amorphous Fe-oxide in vesicles and after py

Description

Massive pyrite." 113 Leached

layer layer layer layer

Chl-smec Chl-smec Chl-smec Chl-smec

Dykes: 110 112 154 156

mixed mixed mixed mixed

Alteration facies*

Sample

Table 2. Description of Pitharokhoma samples

z~ t-" tt~ ~-t

REE MOBILITY IN A MINERALIZED ALTERATION PIPE

159

heterogeneously overprinted by K-feldspar and calcite attributed to lower-temperature alteration (Richards et al. 1989). Similarly, illite in the pipe centre contains expandable smectite layers and is inferred to have formed from hydrothermal sericite in the waning stages of the oreforming episode, or during post-mineralization alteration of the lavas (Richards et al. 1989). Subaerial gossanisation due to the circulation of meteoric waters through the mine has degraded the exposure since exploratory mining in 1983. This supergene oxidation is unlikely to have had any significant effect on the REE inventory of the alteration pipe.

ric acid mixture was evaporated to near dryness and diluted to a known weight with 6 M HC1. The REEs were separated from this solution using a modified cation exchange procedure of Greaves et al. (1989) by a single passage through a small cation exchange column. REE fractions were evaporated to dryness, then diluted to a known weight with 2% HNO3 for analysis. All REE analyses was made on a VG Plasmaquad PQ2 + Inductively Coupled Plasma Mass Spectrometer (ICP-MS) at the Southampton Oceanography Centre.

Sampling and methods

Operating conditions were adjusted for maximum count stability and machine sensitivity using a multi-element solution containing 0.1 ppb of Be, Co, Y, In, La, Re, Bi and Ba. The ion lens system was optimized on J39La. An in-house standard with REEs in chondritic proportions was prepared to facilitate a superior calibration of sample HREEs at low concentrations. Standards were run at the beginning and end of the procedure. Data were acquired in 5 runs of 30 s. Calibration of the results was achieved using appropriate dilutions of the in-house REE standard in a 2% HNO3 matrix. Signal drift was monitored by running a 1 ppb REE standard after every fifth sample and a linear drift correction was applied to the count integrals. Within-run precision was typically better than 3% (2o-). Column recovery was assessed by analyses of solutions spiked with a known quantity of Tm that were loaded onto cation exchange columns, and were typically better than 99%. An off-line custom written computer program, the ICP-MS Data Manipulation Program, was used for data processing (A. Milton pers. comm.). This program applies a drift and a blank correction to the raw data and produces a multi-standard calibration based on the concentration of REEs in the standard solutions. Repeat analyses of the REE fraction of the BHVO-1 international basalt standard obtained using the above procedures were on average accurate to within 4% of published values (Govindaraju 1996) with an external precision averaging 7% (2o-).

Major, trace and rare earth element determinations were made on 18 samples collected from across the eastern alteration pipe of the Pitharokhoma mine (Fig. 3; Table 3). The sample set comprises variably altered and mineralized lavas and dykes, haematized and mineralized interstitial sediments and massive pyrite (Table 2). Before grinding to powders, samples were trimmed to remove surface oxidation. Major and trace elements were determined by X-ray fluorescence (XRF). Major elements were determined on fused beads prepared with 10:1 or 20 : 1 mixtures of Spectroflux 100 and ignited sample powder. Precision of major element data is 1 to 5% (2o-) for SiO2, A1203, Fe203, CaO, K20; 5 to 10% (2o-) for TiO2, MnO, MgO and over 10% (2o-) for Na20 and P205. Trace elements were determined on pressed powder pellets bound with an aqueous solution of polyvinyl alcohol. Precision of trace element data is better than 3 % (2o-) when concentrations are well above detection limit. A modified Compton scatter technique was employed to correct for matrix effects for all wavelengths (Croudace & Gilligan 1990) resulting in detection limits which range from 4 ppm (Nb, Rb, Sr and Zr) to 60 ppm (Ti). X-ray diffraction (XRD) patterns were obtained for powdered samples using an ENRAFN O N I U S PDS120 system with a position sensitive detector (PSD) at The Natural History Museum, London.

Mass spectrometry

Rare earth elements Samples were roasted at 600°C prior to digestion to oxidise any sulphide phases present. Approximately 0.3 g of dried homogenised sample was dissolved in closed Teflon digestion vessels using a combined hydrofluoric acid/perchloric acid digest technique. The hydrofluoric acid/perchlo-

Laser ablation inductively coupled plasmamass spectrometry The REE compositions of secondary alteration phases within the Pitharokhoma samples were determined by laser ablation (LA) ICP-MS with

160

D. M. WELLS E T AL.

a Fisons 'S' operation interface coupled to a frequency quadrupoled N d - Y A G UV Qswitched laser probe operating at 266 nm. REE data were acquired over 15s, with 5s preablation, giving spatially resolved REE analysis from 10-20#m ablation pits. Data were calibrated by comparison with count rates for a synthetically manufactured NIST 612 glass wafer standard reference material (SRM) containing 40 ppm (nominal) of a range of trace elements including the REEs. Differences between the ablation characteristics of the sample and SRM (primarily due to matrix variations) can significantly influence the element concentrations determined. An element present in the sample and the standard at a known concentration can be used as an internal standard for LA ICP-MS analyses. However, an exploratory geochemical survey of the Pitharokhoma samples by LA ICP-MS revealed significant compositional heterogeneity at the < 15#m level. Without internal standardization of the dataset, measured REE concentrations may be subject to an order of magnitude inaccuracy although inter-element ratios will be broadly unaffected. Consequently, the degree of enrichment or depletion of Ce, Eu and the LREEs relative to the HREEs in the chondrite normalized patterns acquired by LA ICP-MS can be interpreted in an equivalent manner to whole rock REE data. The reproducibility of 10 repeat laser shots for the REEs in a NIST SRM is typically 2 to 5% (2a). REE data have been normalized to a chondrite reference material (Evensen et al. 1978). The Ce anomaly, Eu anomaly, and the fractionation between the LREE and the HREE are quantified to facilitate comparison of the sample data. The deviation of Ce and Eu from the other REE can be expressed as:

Ce anomaly = Ce/Ce* = 3Ce,,/(2La,, + Ndn) and Eu anomaly = Eu/Eu* = 2Eu,,/(Sm,, + Gd,,) where the subscript n refers to the normalized values and the superscript * refers to the value obtained by linear interpolation between adjacent elements. In the absence of Gd data for every sample, the chondrite normalized Eu/Sm (Eu,,/Sm,,) ratio has been used to approximate the Eu anomaly. The Ndn/Ybn and Lan/Smn ratios are used to quantify the fractionation of the LREE from the HREE and the relative degree of LREE depletion respectively. All ratios are shown in Table 4.

Results Major and trace element data are presented in Table 3. Geochemical and alteration trends across the eastern pipe are similar to those reported by Richards et al. (1989), summarized in Table 1. Immobile element ratios indicate that the lavas were originally andesites or basaltic andesites with 60 to 55% SiO2 and 2 to 6% MgO (Richards 1987; Richards et al. 1989). Metasomatic changes within each alteration facies were calculated by Richards (1987) by comparing immobile element ratios in the altered lavas to those of available analyses of Troodos volcanic glasses. The smectitic facies is depleted in Si and Ca, and enriched in K, Na, Rb and Ba. These changes can be ascribed mainly to the heterogeneous occurrence of secondary smectite and K-feldspar in these lavas (Gillis & Robinson 1988). Chlorite-smectite mixed layer facies lavas are more Ca-depleted than smectitic facies lavas due to enhanced alteration of Ca-bearing phases. They are variably enriched in Na and Mg, reflecting the degree of plagioclase albitiztion and replacement of the igneous groundmass with mixed-layer chlorite smectite (+discrete chlorite), respectively. The low Na and Sr content of sample 157 reflects the extent of plagioclase replacement by mixed-layer chloritesmectite and chlorite. This facies is generally depleted in K, Rb and Ba, excluding lavas that contain post-mineralization K-feldspar (Richards et al. 1989). Towards the pipe centre, the enhanced albitization of plagioclase (chlorite-albite facies) gives way to chlorite-illite facies alteration where albitized plagioclase is progressively replaced by illite and the igneous groundmass by quartz and Mg-rich chlorite. Leached facies lavas in the centre of the alteration pipe are altered to an illite-quartzpyrite-anatase assemblage. The extent to which fluid flow was focused through the alteration pipe during axial hydrothermal circulation is evident in the magnitude of major and trace element metasomatism of these lavas (Richards et al. 1989). The chemical changes in this facies are due to illitization at the expense of Mg-rich chlorite. The lavas are highly depleted in MgO (up to - 3 . 8 g per 100g; Richards et al. 1989). The elimination of feldspar results in depletions in Ca, Na and Sr with K, Ba and Rb added through illitization (Richards et al. 1989). Lava compositions indicate considerable mobility of A1203 (up to - 10.5 g per 100 g), V and P. Only Ti, Zr and Y are judged to have been immobile

97.0

. 8.9 243 . . 70.8 33.5 242 66.4 17.3 82.9 16.6

Total

As Pb Zn Mo Sb Rb Sr Ba V Co Ga Ni Zr Cr

.

. .

.

5.5 12.4 59 306 55.5 16.4 6.5 68.7 20.3

5 317

99.0

7.6

54.1 1.2 13.6 12.1 0.21 8.2 0.14 0.74 0.71 0.09 0.45

112

.

. .

.

78.2 92.1 390 80.7 17.1 15 45.4 23.9

2240

98.9

50.2 1.0 14.3 13.7 0.34 7.3 2.13 4.1 0.38 0.08 0.35 . 5.0

154

. .

.

343 . . 77.6 326 75.6 16.3 17.7 58.3 29.4

99.0

51.8 1.1 14.9 12.1 0.23 7.3 4.36 3.0 0.13 0.10 0.23 . 3.8

156

.

.

76.9 35.9

94.5 . . 8.1 120 123 198 55.3 17

98.6

3.3

56.8 1.2 14.2 10.1 0.12 3.9 5.08 2.8 0.79 0.12 0.22

146

.

.

24.1 237 . . 12 128 239 236 65 16.7 7.4 71.5 -

97.5

3.7

54.3 1.3 13.2 12.1 0.35 5.2 2.70 2.8 1.52 0.10 0.16

145

.

. .

.

62.9 83.2 368 79.3 16.5 14.9 53.3 20.8

6.2 619

99.2

4.8

50.9 1.2 15.5 13.0 0.15 7.5 0.76 4.6 0.34 0.10 0.30

147

. .

.

15.8 33.2 494 48.1 16.9 8.4 61.7 29.3

8.9 379

99.1

53.2 1.3 18.7 12.7 0.05 2.0 0.14 0.30 0.42 0.11 0.40 0.21 9.5

153

. .

Lavas

4.4 84.4 74.1 267 64.8 14.3 27.1 34.8 35.6

98.8

99.5

53.2 0.81 14.9 9.94 0.26 8.3 2.97 3.41 0.74 0.05 0.55 0.01 4.3

155

M a j o r e l e m e n t s are in wt %; v o l a t i l e s are d e t e r m i n e d as loss on i g n i t i o n ( L O I ) . T r a c e e l e m e n t c o n c e n t r a t i o n s are in p p m . Dash indicates element not detected.

55.9 1.1 13.4 11.1 0.13 5.6 2.83 2.1 0.11 0.11 0.64 . 3.9

SiO2 TiO2 A1203 Fe203 MnO MgO CaO Na20 K20 P205 SO3 CuO LOI

110

Dykes

Table 3. Major and trace element data for Pitharokhoma samples

5.6 5.1 35.5 367 103 16.5 14 61.6 24.7

314

99.4

52.0 1.4 15.9 10.9 0.31 5.5 0.10 . 0.47 0.03 0.94 0.03 11.8

157

.

159 19.5 142 9.4 14.2 16.7 6.5 554 222 140 8.8 27.9 17.5

99.3

1.39 . 0.48 . 15.3

52.8 0.59 5.81 22.5 0.40 0.06

114

.

.

.

26.9 284 20.5 7.3 463 236 648 5.2 16 18.1 28.3

98.5

15

0.10

1.88

54.2 0.42 7.18 19.1 0.35 0.19

143

.

.

.

.

92.1 21.5 237 5.5 12.2 7.7 661 199 447 . 11.8 25.6 19.2

99.1

18.8

0.95 . 0.26

46.6 0.58 3.78 27.8 0.16 0.12

150

.

.

.

.

208 45.2 3640 79.7 11 . . 91.4 334 . . . -

99.4

0.19 . 32.2

0.02

3.88 0.01 0.01 62.9 0.01 0.07

113

.

.

Massive pyrite

33.4 6.30 28.9 . 15 44.5 92.3 109 . . -

99.7

1.1

0.50

-

90.6 0.04 0.82 6.24 0.02 0.26 0.15

Ill

.

.

15 6.70 73.5 . 9 47.9 305 91.2 . 9 . 19.3

99.7

1.1

0.02 0.09

83.6 0.04 1.25 12.8 0.56 0.11

144

.

.

.

.

26.8

25

65.6 33.8 1330

112 25.5 61.8 22.8 -

99.8

11.0

0.04 0.11

0.17 0.03

68.6 0.02 0.49 19.3

149

-

-

4 603

532 78.2 1050 32.4 18.1

99.5

32.3

0.11

5.94 0.03 61.1 0.01 0.05

152

Interstitial sediments

0.0523

1.81

1.48

1t.263

0.681

5.68E-04

3.54E-06 5.13E-06

0.88

3.56

0.262

0.802

Massive pyrite: 113

Interstitial sediments: Ill

144

149

152

*Black smoker fluid

*North Atlantic Seawater ~Pacific Seawater

§Troodos glass 345

§Troodos glass 353

¶ Smectite CY1:86. I

¶ Smectite CY 1:437

10.94

2.39

7.62E-07 4.62E-07

1.32E-03

1.5

0.60

2.5

2.7

0.12

2.3

3.0

1.1

4.6

4.4

1.5

6.6

8. I

9.2

7.3

6.2

5. I

10

Ce

3.71E-07

0.254

0.0995

0.497

0.420

0.0189

0,422

0.515

0.219

0.836

0.760

0.328

1.11

1.36

1.51

1.20

0.987

0.996

1.64

Pr

1.948

0.717

10.41

2.43

3.09E-06 1.68E-06

8.68E-04

1.08

0.441

2.62

1.97

0.0836

2,29

2.56

1.27

5.22

4.24

2.13

6.42

7.66

8.45

6.84

5.51

6.27

9.10

Nd

0.777

0.288

3,85

1.03

6.21E-07 3.25E-07

1.87E-04

0.185

0.0790

0.851

0,52I

0.0158

0,892

0.754

0.534

2.18

1.51

1.16

2.32

2.78

2.79

2.45

1.98

2.64

3.08

Sm

0.214

0.066

1.35

0.419

1.61E-07 8.97E-08

5.42E-04

0.0486

0.0167

0.316

0,170

0.00311

0.214

0,186

0,143

0.953

0.592

0,454

0.828

1.10

1.05

0.948

0.888

0.920

1.03

Eu

5.55

1.63

9.83E-07 5.65E-07

1.70E-04

0.143

0.0700

1.09

0.697

0.0246

1.47

1.23

0.928

2,86

2,09

I. 54

2.98

3.75

3.911

3.52

2.78

3.70

4.18

Gd

Dy

9.44E-08

0.0175

0.011

0.196

0./21

0.00567

0,295

0.267

0.196

0.566

0.412

0.366

0.548

0.701

0.735

0.662

0.519

0.730

0.775

1.38

6.98

2.16

1.03E-06 7.93E-07

1.23E-04

0.0858

0.0610

1.34

0.856

0.0375

2.06

1.88

1.41

4.02

2.92

2_82

3.73

4.87

5.12

4.72

3.63

5.24

5.32

all values p p m

Tb

0.389

0,083

2.52E-07

0.014

0.011

0.29

0_20

0.0074

0.46

0,40

0,31

0.85

0.62

0.60

0.79

1.0

1.0

1.0

0.78

1.1

I.I

Ho

0.384

4.68

1.44

9.15E-07 9.03E-07

4.93E-05

0,034

0.029

0.86

0.62

0.020

1.4

1.2

0.94

2.5

1.8

1.9

2.3

3.0

3.0

2.9

2.3

3.3

3.1

Er

1.48E-07

0.00442

0.00419

0.123

0,0912

0.00242

0.204

0.159

0.141

0.379

0.246

0.321

0.337

0.440

0.414

0.426

0.320

0.490

0.439

Tm

¶ D a t a f r o m Gillis et al. (1992); d a t a represent the range in R E E content of secondary smectite in T r o o d o s lavas.

§ D a t a f r o m Rautcnschlein et al. (1985); data represent the range in R E E content of Troodos volcanic glasses.

$ D a t a from Bau et al. (1996); average of six samples from between 1000-2000 m depth.

* Data from Mitra et al, (1994): black smoker fluid d a t a is recalculated to M g = 0 , white smoker is recalculated to M g = 4 mmol kg -l. Seawater d a t a is for T A G (3400 m),

1.21

I).428

143

0.952

1.28

155

157

114

1.50

153

150

1.92

0,387

147

3.08

2.18

2.52

2.09

154

156

145

1.29

112

Lavas: 146

2.93

Dykes: 110

La

Table 4. REE data (ppm) for Pitharokhoma samples

1.0 0.80 0_91 0.88

0.41 0.14 0.14 0.22

0,66 0,66 0.91 0.93

0.0038 0,0036

4,41E-06 0.96 1.54E-07 0.10

0.024 3.60E-05 9.38E-07

0.91

0,217"0.689

1.44 4.59

1.0

0.048

1.12E-06

0.026

0.79

0.13

0.90

0.98

0.22

0.002

1.06 0.91

0.37

1,0 0.35

1.00

1.1

0.45

1.0

0.31 0.44 0.38

1.1 1.0

0.42

Ce/Ce*

0.52

Lu

0.10 0.62

0.014

1.4

0.97

0.91

2.6

1.6

2.5

2.2

2.9

2.6

2.8

2.1

3.3

2.8

Yb

0.89

0.98

0.63

0.63

9.1

0.88

0.67

1.0

0.86

0.48

0.57

0.59

0.62

1.2

1.0

1.0

0.96

1,0

0.97

0,99

1.2

0.90

0.88

Eu/Eu*

0,60

0.58 0.57 0.65

0.73

0.54

9.9

3.6

1.9

2.3

2,1

1.1

2.2

2.1

0.67

1.0

0.50

0.37

0.63

0_21

0.52

0.57

0.70

0.56

0.66

0.31

0.93 0.61

1.1

0,73

0,69

7.7

0.70

0.56

0.99

0.86

0.52

0.64

0.65

0.71

1.2

1.0

1.0

0,95

1.0

1.0

1.0

1,2

0.93

0,89

0,79

0,59

0.53

1,1

8.4

15

6.0

1.2

I. 1

2. I

0.59

0.92

0.49

0.71

0.95

0.30

1.00

0.91

1.1

0.85

0.92

0.66

1.1

Eu,,/Sm,, La,,/Sm,, Nd,]Yb,,

REE MOBILITY IN A MINERALIZED ALTERATION PIPE

163

100

~ ' ~ . .' . ' - -~. : . ~-: : : :~ ~ o~

.

::::::::::::::::::::::::::::: : ~!::::,i: ::'-: ~. ~:.:~,.~.::;_

~

10

E o

'~

" - ~ 112 ~ ~

1

o

154 156

Dykes

Chlorite-smectite mixed layer facies L

0.1

I

~

La Ce Pr Nd

I

F

J

J

I

I

I

~

~

P

J

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Fig. 4. Chondrite normalized whole rock REE data for Pitharokhoma dykes. The shaded area shows the range in REE content of Troodos volcanic glasses. Data are from Rautenschlein et al. (1985).

during axial alteration (Richards et al. 1989). Similar trends are noted for leached facies samples analysed in this study. For example sample 150 has a TiOz/A1203 ratio of 0.15, compared with values typically less than 0.1 for most Troodos Ti-rich lavas (Robinson et al. 1983). Whole rock REE data are shown in Table 4. Chondrite normalized REE data (Figs 4 & 5a, b) are plotted with a field representing the range of REE compositions of Troodos volcanic glasses analysed by Rautenschlein et al. (1985). These glasses are presumed to be free from the effects of secondary alteration and have a R E E composition similar to N-MORB (Fig. 1).

P i t h a r o k h o m a dykes Approximately 20% of the rocks exposed in the pit are dykes. Some of these were probably intruded after the peak of hydrothermal activity, as they are differentially altered with respect to their host lavas (Richards 1987). The alteration of samples 110, 154 and 156 is equivalent to adjacent lavas, hence their intrusion is inferred to pre-date axial hydrothermal alteration. They are altered to chlorite-smectite mixed layer facies assemblages, but are less albitized and chloritized than the lavas they intrude (e.g. 155) This implies the dykes were less susceptible to hydrothermal alteration than the enclosing lavas. The most altered dyke analysed (112)

intrudes illitized lavas at the pipe centre (e.g. 114), but contains secondary chlorite and mixed layer chlorite-smectites rather than illite. Chondrite normalized whole rock REE patterns for dyke samples are shown in Fig. 4. Absolute REE concentrations (5.51-9.10 ppm Nd) are comparable to Troodos glasses (2.4310.41 ppm Nd). They exhibit LREE depletions and Eu anomalies (Lan/Smn = 0.31-0.66; Eu/ Eu* = 0.88-1.2) that are slightly smaller or larger than those of Troodos glasses (Lan/ Sin,, = 0.54-0.62; Eu/Eu* = 0.89-1.0). Hydrothermal alteration of these lavas has not caused any significant net REE mobilization. There is no correlation between the degree of LREE depletion and the shape or magnitude of the Eu anomaly.

Pitharokhoma lavas Smectitic and chlorite-smectite mixed layer facies lavas at the periphery of the alteration pipe (Fig. 5a) and leached facies lavas at the pipe centre (Fig. 5b) display contrasting REE patterns.

Smectitic and chlorite-smectite mixed la}~er facies. Chondrite normalized REE patterns for lavas from the smectitic and chlorite-smectite mixed layer alteration facies generally parallel Troodos glasses with some variability in the LREEs (Fig. 5a). There is little evidence for net

D. M. WELLS ET AL.

164 100

10

I-.-145 +1461 ~147 / ---n-- 153/ --o- 155j -°-157]

Laras

Smectitic & chlorite-smectite mixed layer facies 0.1

f

~l

La Ce Pr Nd

(a)

1

I

i

i

r

i

i

i

L

I

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

100

10 N

I i

+

143 150

Lavas

114

Leached facies 0.1

(b)

_ _

i

La Ce

~

i

i

Pr Nd

Sm Eu

f

Gd Tb Dy

~

Ho

r

~

~

r

Er Tm Yb Lu

Fig. 5. Chondrite normalized whole rock REE data for (a) Smectitic and chlorite-smectite mixed layer facies lavas (b) Leached facies lavas. The shaded areas in (a) and (b) show the range in REE content of Troodos volcanic glasses (data are from Rautenschlein et al. 1985).

REE mobilization in smectitic (146) and most of the chlorite-smectite mixed layer facies lavas (145, 147, 155). They have REE concentrations (4.24-8.45 ppm Nd) and ratios (Lan/Smn = 0.520.63; Eu/Eu* = 0.96-1.0) similar to Troodos glasses (2.43-10.41 ppm Nd; La,/Smn = 0.540.62; Eu/Eu* = 0.89-1.0). Two samples from the chlorite-smectite mixed layer facies (153, 157) are more L R E E depleted (La,,/Smn = 0.21 and 0.37) than Troodos glasses (Lan /Smn = 0.540.62) indicating a net mobilization of LREEs during crustal alteration. The R E E content of secondary alteration

phases in some of these lavas was determined by LA ICP-MS. The minerals analysed were identified petrographically and/or by X R D analyses. Sample 146 from the periphery of the mine is altered to a smectitic alteration assemblage that is inferred to post-date axial mineralization (Richards et al. 1989). T w i n n e d feldspars in this lava are essentially unaltered. Smectite is the principal secondary phase and occurs replacing the mesostasis and intergrown with fracture filling zeolites. LA ICP-MS analyses show void filling smectite to be slightly L R E E enriched (La,/Smn = 0.44-0.58) and Eu

REE MOBILITY IN A MINERALIZED ALTERATION PIPE

I

1000

165

Smectiticfacies

1

............ i .2", i

0.1 La Ce

Pr Nd

i

_t

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Matrix filling smectite/amorphous Fe-oxide (7) Zeolites (mordenite/cowlesite; 3) Void filling smectite (2) :-

146 (whole rock)

(c)

ititemixedlayerfacies

1000

100

-~

10

i~

1

Fracture filling amorphous Fe-oxide/goethite (21)

d

Chlorite-smectite(4) 0.1

Matrix filling Fe-oxide/chlorite-smectite (3) 153 (whole rock)

0.01

I

~_ ....

L

La Ce Pr Nd

I

i

I

~

.....

J

~

l

.1____

i

I

I

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

(d) Fig. 5. Chondrite normalized LA ICP-MS REE data for alteration phases in samples; (c) 146 from smectitic facies lavas. The dashed line shows the range in REE content of secondary smectite in Troodos lavas (Gillis 1992) (d) 153 from chlorite smectite mixed layer facies lavas

etal.

depleted (Eun/Smn = 0.52 and 0.55) relative to the whole rock (Fig. 5c). lntergrown zeolites (mordenite and cowlesite) which fill vesicles and fractures are L R E E enriched compared to the whole rock (Lan/Smn = 1.9-3.8 cf. 0.70) and variably Eu depleted (Eu~/Sm, = 0.14-0.74 cf.

1.0). Intimately associated smectite and amorphous Fe-oxides which replace much of the igneous groundmass are typically more L R E E and Eu depleted (Lan/Sm, = 0.42-0.72; Eu/ Eu* = 0.64-0.84) than the whole rock (La,,/ Sm, = 0.70; Eu/Eu* = 0.97).

166

D. M. WELLS E T AL. I000

Chlorite-smectite mixed layer facies

100

!

lO Matrix filling Fe-oxide (3)

;5 ~

Chlorite-smectite (2)

1

145 (whole rock)

0.1

I

J-.-

I

I

La Ce Pr Nd

(e)

I

i

I

_.k__J

I

I

I

I

I

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

looo F

Leached facies

100

10

1

0,1

(0

I

L___

Illite (3) * 114 (whole rock) i

J

I

La Ce Pr Nd

__~

1

t

I

j

F

I

1

~

_

I

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Fig. 5. (e) 145 from chlorite smectite mixed layer facies lavas (f) 114 from leached facies lavas. The numbers in parentheses indicate the number of analyses represented by each compositional field.

Sample 153 from the chlorite-smectite mixed layer facies contains ~ 4 0 % secondary chloritesmectite developed primarily within the igneous groundmass and ~ 1 0 % matrix- and fracturefilling amorphous Fe-oxides with minor goethite (Table 2). LA I C P - M S R E E p a t t e r n s of fracture-filling amorphous Fe-oxides are relatively flat to H R E E enriched and generally more L R E E depleted (Lan/Smn = 0.09-0.53) than the whole rock (Lan/Sm,, = 0.21; Fig. 5d). Poorly crystalline Fe-oxide stained mixed layer chloritesmectites replace both igneous feldspars and groundmass adjacent to amorphous Fe-oxide filled fractures. They display flat LREEs (Lan/

Smn = 0.67-1.0) and some H R E E enrichment (Ndn/Yb,, = 0.25-0.28) compared to the whole rock (Nd,,/Ybn = 0.30). Chlorite-smectites without overprinting amorphous Fe-oxides have REE patterns that are L R E E enriched with a larger negative Eu anomaly (Lan/Smn = 2.7-4.4; Eu/Eu* ~ 0.62-0.79) than the whole rock pattern (Lan/Smn = 0.21; Eu/Eu * = 1.0). The LA ICP-MS data reveal that the REE composition of amorphous Fe-oxides that overprint the chloritic alteration is dominating the whole rock pattern of this lava, due to their elevated REE content. Sample 145 from the chlorite-smectite mixed

REE MOBILITY IN A MINERALIZED ALTERATION PIPE layer facies comprises lava fragments partially replaced by chlorite-smectite within a matrix of amorphous Fe-oxides, quartz and chloritesmectite. Matrix-filling amorphous Fe-oxides in this sample have REE compositions that mimic pristine basalt (Fig. 5e). Chlorite-smectites show somewhat flatter REE patterns (Lan/Smn = 1.01.5) than in sample 153 (Lan/Smn = 2.7-4.4). Leached facies. Lavas from the leached facies (114, 143, 150) comprise chiefly illite, quartz and pyrite. Consequently, the whole rock R E E patterns of these lavas reflect the REE composition and modal abundance of these secondary phases. They have whole rock REE contents that are similar to, or slightly lower than Troodos glasses (1.27-2.56 ppm cf. 2.4310.41 ppm Nd) and are similarly LREE depleted (Lan/Sm, = 0.50-1.0 cf. 0.54-0.62; Fig. 5b). Metasomatic gain of SiO2 and FeS2 during axial alteration manifested as replacive quartz and pyrite (Richards et al. 1989) will tend to dilute the whole rock REE concentrations and account for a degree of the observed depletion. Eu is significantly more depleted than the other REEs (Eu/Eu* = 0.57~0.62 cf. 0.89-1.0 for Troodos glasses). Quartz and pyrite in sample 114 have REE contents below the detection limit of the LA ICP-MS technique. Thus, illite is inferred to be the major repository for REEs in the centre of the alteration pipe. Fine-grained illite in sample 114 occurs in a granular matrix of quartz and pyrite with rare preservation of igneous textures, and shows a range of REE compositions (Fig. 5 f). H R E E enriched illite (Ndn/Yb~ -- 0.19) with a Eu depletion comparable to the whole rock (Eu/Eu* = 0.56 cf. 0.57) can be inferred to dominate the whole rock REE composition. Another composition is slightly more H R E E enriched (Nd~/Yb,, = 0.47) than the whole rock without any Eu enrichment or depletion (Eu/ Eu* = 1.0). A third illite composition is extremely L R E E depleted (La,/Smn = 0.02) and H R E E enriched (Nd~/Ybn = 0.32) relative to the whole rock, with a REE composition similar to low temperature amorphous Fe-oxides in sample 153 (Fig. 5d). Interstitial sediments. The whole rock REE data

for interstitial sediments within the alteration pipe defines two groups with contrasting REE characteristics (Fig. 6a); haematitic jasper from the chlorite-smectite mixed layer facies (111, 144) and highly mineralized sediments from the leached facies at the pipe centre (149, 152). Interstitial haematitic jasper has a relatively flat REE pattern (Nd~/Yb,, = 1.1 and 1.2) with a

167

p r o n o u n c e d n e g a t i v e Ce a n o m a l y (Ce/ Ce* = 0.66), demonstrating the influence of seawater REEs (Fig. 1). Haematitic jasper (111) has a REE composition that is slightly more H R E E enriched and Ce depleted (Ce/ Ce* = 0.46-0.50; Nd,,/Yb,, = 0.58-1.1) than the whole rock (Ce/Ce* : 0.66; Nd,JYbn = 1.1; Fig. 6 b). Void filling smectite in this sample is LREE enriched (Lan/Smn = 4.0-5.7) compared with the whole rock (Lan/Sm. = 2.2) with no Ce depletion (Fig. 6b). Interstitial sediments from the pipe centre are intensely altered to granular aggregates of pyrite and quartz with minor illite, and account for a large proportion of the total sulphide mineralization within the alteration pipe (Richards et al. 1989). They are depleted in REEs (Nd = 1.08 and 0.441 ppm) and show a striking LREE enrichment (Ndn/Yb,~ = 6.0 and 15) relative to hematite jasper (Nd = 1.97 and 2.62 ppm; Nd,,/ Ybn = 1.1 and 1.2). Quartz and pyrite in sample 149 have REEs below the detection limit of LA ICP-MS. A compositionally insignificant phase was identified by SEM as a K-A1 silicate, inferred to be illite or possibly hydrothermal sericite (Richards et al. 1989). The REE content of the illite/sericite evidently controls the whole rock REE composition, despite its low modal abundance (less than 5%). It contains abundant REEs (Fig. 6c) and is more enriched in LREEs (La,~/Sm~ = 3.5-7.5) and depleted in Eu (Eu/Eu* = 0.4563) than the whole rock (La,/Sm, = 2.1; Eu/Eu* = 0.67). While apatite has been described as a minor phase in leached facies interstitial sediments (Richards 1987; Richards et al. 1989) it was not observed in these samples, consistent with low measured P205 contents (below detection limit). Massive pyritic mineralization. Massive sulphide

in the centre of the alteration pipe consists of monomineralic pyrite (113) with very low levels of REEs (Nd = 0.0836 ppm). It is L R E E enriched (Ndn/Ybn = 2.1) but depleted in the middle REEs, particularly Eu (Eu/Eu* = 0.48; Fig. 6a).

Discussion The chondrite normalized REE patterns for North Atlantic and Pacific seawater (Mitra et al. 1994; Bau et al. 1996), a typical black smoker fluid (Mitra et al. 1994) and the range of REE compositions of Troodos volcanic glasses (Rautenschlein et al. 1985) are shown in Fig. 1. The REE compositions of North Atlantic and Pacific seawater are characterized by a pronounced

168

ET AL.

D . M . WELLS

lO ~

C h l o r i t e - s m e c t i t e m i x e d layer f a c i e s

1

0.1

--a-144 149 001.

~

La

-~

Ce

~

Pr

~

Nd

---o-111 x 113

~

'

'

Sm

Eu

Gd

¢. 152 1

/

~

'

Tb

'

'

Dy Ho

- - ~ ~ '

Er

'

Tm Yb

Lu

(a) 100

Chlorite-smectite mixed layer facies

10 "N

0°1

_

La

_

~

Ce

_

Pr

i

Nd

_

~

Hematitic jasper (3)

~

Smectite (2)

~

111 (whole rock) L

'

J

r

Sm Eu Od

i

__

L

~

Tb Dy Ho

Er

I

I

Tm Yb

]

Lu

(b) 100

Leached facies

10

.=

0.1

Illite/sericite (9) --II--

149 (whole rock)

0.01 La (c)

Ce

Pr

Nd

Sm Eu

Gd

Tb

Dy Ho

Er

Tm Yb

Lu

REE MOBILITY IN A MINERALIZED ALTERATION PIPE negative Ce anomaly and H R E E enrichment. Similar REE characteristics would be predicted for Cretaceous seawater, given that the REE pattern of seawater is governed by the oxidation state of the ocean and the physical properties of the REEs. Black smoker fluids are enriched in REEs (1 to 3 orders of magnitude) over seawater concentrations. Chondrite-normalized black smoker REE patterns show a large positive Eu anomaly, LREE enrichment but no Ce anomaly. This pattern is common to all known hightemperature fluids from sediment-free ridges, although there is a some variation in the size of the Eu anomaly between different vent sites (Klinkhammer et al. 1994; Mitra et al. 1994). Lavas and dykes from the alteration pipe show either no change in REE content or a depletion in the LREE ± Eu. Quantification of REE mobility in the Pitharokhoma samples is problematic because all lavas and dykes analysed have been subject to a degree of hydrothermal alteration. Immobile element ratios indicate significant primary geochemical variability within the Pitharokhoma lavas (Richards 1987) hence the REE content of a pristine lava or dyke cannot be assessed with confidence.

169

5c,d). It is inferred that any high-temperature (LREE enriched) signature has been overprinted by REEs sourced from a lower-temperature hydrothermal fluid. This supports the inference of Richards et al. (1989) that illite formed from non-expandable sericite during the waning stages of axial alteration, or during off-axis alteration at temperatures less than 260°C (Richards et al. 1989). The different illite REE compositions in sample 114 (Fig. 5f) may indicate variable degrees of crystallographic control on REE uptake from the fluid, or differences in fluid composition arising from retrograde precipitation-dissolution reactions within the lavas. The LA ICP-MS REE data indicate that the low-temperature phyllosilicates are the major repository for REEs in leached facies lavas in the core of the alteration pipe. LREE enriched illite/ sericite in the mineralized interstitial sediments with a REE composition derived from ~350°C axial hydrothermal fluids was apparently stable during lower-temperature alteration, conceivably due to its occurrence within a robust quartzpyrite assemblage. C h l o r i t e - s m e c t i t e m i x e d layer f a c i e s

Leached facies

Secondary illite is the primary repository for REEs in the leached facies lavas and interstitial sediments at the centre of the alteration pipe. Illite in the lavas (e.g. 114; Fig. 5f) and illite/ sericite in the mineralized sediments (e.g. 149; Fig. 6c) show contrasting REE patterns, that are interpreted to reflect the REE signatures of fluids they have been altered by, or precipitated from. It is inferred that LREE enriched pattern of illite/sericite in the mineralized interstitial sediments was acquired from upwelling LREE enriched hydrothermal fluids, akin to modern black smoker fluids (Fig. 1). This interpretation is consistent with extreme Mg depletion of the lavas and inferred alteration temperatures of 300-370°C within the pipe centre (Jensenius 1984; Richards 1987). In contrast, illite in the lavas that enclose these sediments has REE compositions similar to amorphous Fe-oxides and intergrown amorphous Fe-oxides/smectite in chlorite-smectite mixed layer and smectitic facies lavas (Fig.

Mixed layer chlorite-smectites in the peripheral alteration facies of the pipe are inferred to have formed at temperatures of ~200°C by analogy with Icelandic geotherma! systems (Richards et al. 1989). Temperatures of formation of 80 to 240°C have been inferred for chlorite in upper dyke zone of the Troodos ophiolite (Gillis & Robinson 1990b). The alteration mineralogy of these lavas reflects mixing of upwelling black smoker fluids with relatively unreacted seawater at the periphery of the axial upflow zone (Richards et al. 1989). The variably LREE enriched composition of chlorite and mixed layer chlorite-smectites in Pitharokhoma lavas (Lan/Smn = 1.0-4.4; Fig. 5d~e) is inferred to reflect precipitation from mixtures of entrained seawater and h y d r o t h e r m a l fluid, that are dominated by hydrothermal fluid rather than seawater REEs (Fig. 1). Amorphous Fe-oxides that overprint chloritic alteration in these lavas are variably LREE depleted (Fig. 5d,e). These patterns are inferred to reflect the REE content of low-temperature

Fig. 6. (a) Chondrite normalized whole rock REE data for interstitial sediments from the chlorite-smectite mixed layer and leached facies and massive pyrite from the leached facies (113). Chondrite normalized LA ICP-MS REE data for alteration phases in interstitial sediments in samples (b) 111 from chlorite-smectite mixed layer facies lavas (e) 149 from leached facies lavas. Numbers in parentheses indicate the number of analyses for each field.

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hydrothermal fluids they were altered, as Feoxides have the capacity to scavenge REEs from solution without significant fractionation (Koeppenkastrop & De Carlo 1992). Large volumes of seawater circulate freely through the permeable volcanic section of oceanic lithosphere (Alt 1995). Because basalt has a greater REE content than seawater (Fig. 1), even minor mobilization of REEs during this circulation would produce solutions with rock-dominated REE compositions. The LA ICP-MS REE data show that the seawater-derived fluids responsible for low-temperature alteration were variably LREE depleted with a REE composition that mimics pristine basalt, as proposed by Gillis et al. (1992). Dyke samples analysed in this study are altered to chlorite-smectite mixed layer facies assemblages, and excluding one sample (112) are hosted by similarly altered lavas. Sample 112 contains secondary chlorite and mixed layer chlorite-smectites and intrudes illitized lavas in the pipe centre. It is inferred that this dyke was relatively impermeable to ascending hydrothermal fluids (resulting in the incomplete elimination of Mg-bearing phases), or was intruded following the peak of hydrothermal alteration. In the latter case, a degree of seawater entrainment into the axial upflow zone is implied by the presence of chlorite in the otherwise Mg depleted pipe centre. The REE patterns of interstitial haematized jasper in chlorite-smectite mixed layer facies lavas (Fig. 6a) reflect uptake of REEs from seawater (Fig. 1) via scavenging mechanisms prior to, or following sedimentation on the Troodos seafloor, and are similar to ridge crest metalliferous sediments (Owen & Olivarez 1988; German et al. 1990) and Troodos umbers (Robertson & Fleet 1976). Smectitic facies

Zeolite and smectite precipitation in voids is characteristic of low-temperature ( < 5 0 to 100°C) alteration of Troodos lavas (Gillis & Robinson 1990a,b). The LREE enriched composition of void-filling smectite in smectitic facies Pitharokhoma lavas is comparable to slnectites in Troodos lavas analysed by Gillis et al. (1992), shown in Fig. 5 c. However, void filling smectite in haematized jasper from the chlorite-smectite mixed layer facies is LREE enriched (Fig. 6 b). Similarly, the REE compositions of mixed layer chlorite smectites, amorphous Fe-oxides and illite vary between samples (Figs 5c-f; 6b,c). Thus, the REE content of a Pitharokhoma lava or interpillow sediment

reflects the REE composition of hydrothermal fluids that have altered them at both high and low-temperatures, and is not a simple function of modal alteration mineralogy. Intergrown void-filling mordenite and cowlesite show a LREE enrichment and a striking Eu depletion, whereas void-filling smectite in the same sample is LREE depleted (Fig. 5c). Assuming the low-temperature fluid had a REE content depicted by amorphous Fe-oxides (Fig. 5d,e) the zeolites have preferentially incorporated LREEs and discriminated against Eu. R E E c o m p o s i t i o n o f the m i n e r a l i z i n g f l u i d

Hydrothermal fluids venting at the seafloor at mid-ocean ridges share a characteristic LREE and Eu enrichment (e.g. Klinkhammer et al. 1994; Fig. 1). Lavas and sediments in the centre of the alteration pipe were altered by ancient analogues of black smoker fluids in a zone of axial hydrothermal alteration (Jensenius 1984; Richards et al. 1989). Secondary chlorite and chlorit~smectites that precipitated at the periphery of the upflow zone possess a LREE enrichment, that is inferred to reflect the REE composition of the hydrothermal fluid. However, none of the secondary phases analysed possess a striking Eu enrichment, although a small positive Eu anomaly has been noted in the REE patterns of chlorite from the sheeted dyke complex (Gillis et al. 1992). Pyrite from the pipe centre is similarly LREE enriched and lacks a positive Eu anomaly (113, Fig. 6a). REE substitution into sulphide phases from the TAG sulphide mound and other oceanic deposits appears to be strongly influenced by crystallographic controls (Morgan & Wandless 1980; Alt 1988; Barrett et al. 1990; Mills & Elderfield 1995). Smaller HREE cations are more easily accommodated in the sulphide lattice than the larger LREEs. Therefore, the REE composition of sulphides reflects that of the parental hydrothermal fluid, albeit with a smaller positive Eu anomaly and a relative HREE enrichment (Mills & Elderfield 1995). Given that the partitioning of REEs between hydrothermal fluid and sulphide is relatively well understood (Morgan & Wandless 1980), the REE composition of pyrite in the Pitharokhoma alteration pipe suggests that the parental hydrothermal fluid had a higher Ndn/Yb,, ratio than the pyrite, and lacked the striking Eu enrichment that is characteristic of seafloor vent fluids. Preferential mobilization of Eu in the divalent state has been invoked to explain the Eu enrichments of vent fluids (Sverjensky 1984;

REE MOBILITY IN A MINERALIZED ALTERATION PIPE Wood 1990; Bau 1991). Divalent Eu dominates at temperatures in excess of 250°C, while at lower temperatures the relative stability of divalent and trivalent Eu will also depend on pH and complexing ligands in solution (Sverjensky 1984). The absence of a Eu enrichment in the pyrite may reflect subsurface processes operating within the upftow zone relating to the temperature, pressure, Eh and pH of the mineralizing system. This has been implied for the Snake Pit vent field on the Mid-Atlantic Ridge, where venting fluids are characteristically Eu enriched, but sulphides in hydrothermal sediments do not show a Eu anomaly, and are inferred to be precipitating from fluids with a different composition to those exiting at the seafloor (Gillis et al. 1990). Alternatively, Troodos hydrothermal fluids may have differed in some fundamental respect to modern vent fluids (relating to redox conditions in the subsurface of the hydrothermal system) which is reflected in the REE composition of the alteration minerals. Axial

alteration

versus

crustal

ageing

processes

The LA ICP-MS REE data demonstrate that during high- and low-temperature hydrothermal alteration there may have been some relative loss or gain of LREEs and HREEs, that did not necessarily cause any substantial net REE mobilization. Although leached facies lavas in the pipe centre were extensively altered by ~350°C hydrothermal fluids, their REE composition is dominated by retrogressive illite and records interaction with a low-temperature fluid, interstitial haematized jasper at the periphery of the pipe has a similar REE composition to ridge crest metalliferous sediments and reflects seafloor sedimentation processes. Mineralized sediments have REE compositions that are inferred to reflect axial alteration by ~350°C hydrotherreal fluids. The REE composition of chlorite-smectite mixed layer facies lavas comprises a component of the primary REE inventory (due to the incomplete alteration of igneous phases), with some high-temperature (mixed layer chloritesmectites + chlorite) and low-temperature REEs (amorphous Fe-oxides). The elevated REE content of overprinting amorphous Fe-oxides controls the whole rock LREE depletion of some lavas and dykes from the chlorite-smectite mixed layer facies (e.g. 153, 157, 112; Figs 4 & 5a) although they also contain variably LREE

171

enriched mixed layer c h l o r i t e - s m e c t i t e s + chlorite. The REE content of smectitic facies lavas at the periphery of the alteration pipe is chiefly a primary inventory, with low-temperature smectite and amorphous Fe-oxides which have REE patterns that mimic fresh basalt. Low-temperature alteration phases have been identified as the major repository for the REEs in many of the altered lavas. Lavas with net depletions in the LREE + Eu have whole rock REE compositions dominated by low-temperature alteration phases (e.g. 153, Fig. 5a; 114,143, 150; Fig. 5b). These data suggest that much of the REE signature of the alteration pipe reflects off-axis low-temperature crustal alteration rather than axial alteration by mineralizing hydrothermal fluids. R E E m o b i l i t y in h y d r o t h e r m a l s y s t e m s

REE depletions (~10 to 1200 ppm of the LREEs) are frequently observed in the alteration zones beneath ancient volcanic-hosted massive sulphide deposits, and often attributed to interaction with ore-forming fluids in the subsurface of the hydrothermal system (Graf 1977; Baker & de Groot 1983; Campbell et al. 1984; Bence & Taylor 1985; MacLean 1988; Whitford et al. 1988; MacLean & Hoy 1991; Regba et al. 1991). Less frequently, REE mobility has been implied to post-date ore formation (Schandl & Gorton 1991), or magmatic fluids have been invoked to account for some of the observed depletions (Schade et al. 1989). A negligible capacity for hydrothermal fluids to mobilize REEs during hydrothermal circulation is inferred from the ubiquitously low REE inventories of seafloor vent fluids (Michard et al. 1983; Campbell et al. 1988; Mitra et al. 1994; Table 4). Studies of hydrothermal systems at mid-oceanic ridges predict that the REE budget of oceanic crust is unmodified during alteration at inferred water/rock ratios of ~ 1 (Michard et al. 1983; Michard & Albar6de 1986; Bau 1991; Klinkhammer et al. 1994). In addition, thermodynamic considerations suggest the observed REE depletions beneath ancient massive sulphide deposits are unlikely to have been achieved by ancient analogues of black smoker fluids, even for systems with water/rock ratios of ~1000 (Wood & Williams-Jones 1994). The extreme REE mobility identified in the altered zones beneath some massive sulphide deposits (e.g. Baker & de Groot 1983; Campbell et al. 1984; Bence & Taylor 1985; MacLean & Hoy 1991; Schandl & Gorton 1991) is not evident in the stockwork mineralized lavas of the Pithar-

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okhoma alteration pipe. Determination of the REE mineralogy of stockwork lavas by an in situ method such as LA ICP-MS might clarify the significance of the extreme REE mobility observed in many other alteration pipes. The inferred hydrothermal circulation for Troodos (Bickle & Teagle 1992; Bickle et al. this volume) is considerably greater than for typical mid-ocean ridges (Morton & Sleep 1985). Hydrothermal alteration at water/rock ratios of ~20 to 1000 has been inferred for the formation of epidosites (quartz-epidote sphene assemblage) near the base of the sheeted dyke complex, that are considered to represent deep upflow zones of Mg-depleted hydrothermal fluids (Richardson et al. 1987; Seyfried et al. 1988; Schiffman et al. 1990). Similarly, it is evident from the extreme metasomatism of Pitharokhoma leached facies lavas that large volumes of ~350°C hydrothermal fluid were focused through this fossilized upflow zone (Richards et al. 1989). While studies of active vent systems predict little mobilization of REEs during axial hydrothermal circulation (e.g. Klinkhammer et al. 1994; Bau 1991; Michard et al. 1983; Michard & Albar~de 1986) these systems are in an early (high-temperature) stage of evolution in comparison to the mature hydrothermal systems preserved within Troodos as superimposed alteration episodes. It is not possible to assess the extent of REE mobilization from the Pitharokhoma lavas by ancient analogues of black smoker fluids during axial hydrothermal alteration, as the REE inventory now observed is chiefly a low-temperature overprint, associated with ageing of the oceanic basement.

Conclusions The major and trace element geochemistry of the lavas of the eastern alteration pipe of the Pitharokhoma deposit described by Richards et al. (1989) reflects increasing hydrothermal alteration towards the centre of a zone of focused hydrothermal fluid upflow within Troodos oceanic lithosphere, with heterogeneous overprinting by lower-temperature phases. LA ICP-MS REE data for stockwork mineralized lavas, dykes and interstitial sediments from the eastern alteration pipe indicate that REE mobility was associated with the development of both high- (~200 to 350°C) and lowtemperature (< 100°C) alteration phases that precipitated within contrasting alteration regimes (discharge- and recharge-dominated, respectively). Lavas and sediments in the centre of the alteration pipe were altered by ancient

analogues of black smoker fluids upwelling in a zone of axial hydrothermal alteration (Jensenius 1984; Richards et al. 1989). The LREE enriched composition of pyrite, chlorite and chlorite-smectite is inferred to reflect the composition of the ~350°C hydrothermal fluid and contrasts with the typically LREE depleted patterns of lower-temperature amorphous Feoxides and smectite. The whole rock REE compositions of altered lavas and dykes thus reflect the REE composition of hydrothermal fluids that have altered them at both high- and low-temperatures. During hydrothermal alteration there may have been some relative loss or gain of LREEs and HREEs, that did not induce significant net REE mobilization in most of the analysed lavas and dykes. Low-temperature alteration phases are the major repository for the REEs in lavas that show LREE 4-Eu depletions relative to pristine volcanic glass compositions. Hence much of the REE signature of the alteration pipe reflects the protracted alteration of the oceanic basement which continued for ,-~20Ma following crustal accretion (Staudigel et al. 1986; Staudigel & Gillis 1990) rather than axial hydrothermal alteration. This study has demonstrated the successful application of LA ICP-MS in identifying alteration phases which host REEs within complex matrices, although as of yet there is no accepted calibration protocol for the method (Norman et al. 1996). C. Xenophontos at the Cyprus Geological Survey Department and D. Teagle are thanked for their contributions to a successful field season. J. Cann is thanked for his assistance in the field, particularly for an introduction to the geology of the Pitharokhoma Mine. B. Nesbitt provided laboratory facilities for ICP-MS preparation. A. Milton is thanked for assistance with ICP-MS analyses and I. Croudace for XRF analyses. G. Cressey is thanked for XRD analyses performed at the Natural History Museum. This manuscript was greatly improved following comments from two anonymous reviewers. This research was funded by NERC grant GR9/01983.

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REGBA, M., AGRINIER, P. PFLUMIO, C. • LOUBET, M. 1991. A geochemical study of a fossil oceanic discharge zone in the Oman ophiolite (Zuha sulphide prospect): evidence for a polyphased hydrothermal history. In: PETERS,T., NICOLAS,A. ~; COLEMAN, R. G. (eds) Ophiolite Genesis' and Evolution of the Oceanic Lithosphere. Ministry of Petroleum and Minerals, Sultanate of Oman, 353383. RICHARDS, H. G. 1987. Petrology and Geochemistry of Hydrothermal Alteration Pipes in the Troodos Ophiolite, Cyprus. Unpublished PhD thesis, University of Newcastle-upon-Tyne. - & BOYLE, J. F. 1986. Origin, alteration and mineralization of the inter-lava metalliferous sediments of the Troodos Ophiolite, Cyprus. In. GALLAGHER,M. J., IXER, R. A., NEARY, C. R. & PRITCHARD, H. M. (eds) Metallogeny of Basic and Ultrabasic Rocks. The Institution of Mining and Metallogeny, 21-31. , CANN, J. R. & JENSENUS, J. 1989. Mineralogical zonation and metasomatism of the alteration pipes of Cyprus sulphide deposits. Economic Geology, 84, 91-115. RICHARDSON, C. J., CANN, J. R., RICHARDS, H. G. & COWAN,J. G. 1987. Metal-depleted root zones of the Troodos ore-forming hydrothermal systems, Cyprus. Earth and Planetary Science Letters, 84, 243-253. ROBERTSON, A. H. F. & FLEET, A. J. 1976. The origins of rare earths in metalliferous sediments of the Troodos Massif, Cyprus. Earth and Planetary Science Letters, 28, 385-394. ROBINSON, P. T., MELSON, W. G., O'HEARN, T. & SCHMINCKE, H.-U. 1983. Volcanic glass compositions of the Troodos ophiolite, Cyprus. Geology, 11, 400~404. SACCOOA, P. & GILLIS, K. 1995. Hydrothermal upflow zones in the oceanic crust. Earth and Planetary Science Letters, 136, 1-16. SCHADE, J., CORNELL, D. H. & THEART, H. F. 1989. Rare earth element and isotopic evidence for the genesis of the Prieska massive sulphide deposit, South Africa. Economic Geology, 84, 49-63. SCHANDL, E. S. & GORTON, M. P. 1991. Postore mobilization of the rare earth elements at Kidd Creek and other Archean massive sulphide deposits. Economic Geology, 86, 1546-1553. SCHIEFMAN, P. & SMITH, B. M. 1988. Petrology and oxygen isotope geochemistry of a fossil seawater hydrothermal system within the Solea graben, northern Troodos ophiolite, Cyprus. Journal of Geophysical Research, 93, 4612~4624. - - , BETTISON, L. A. & SMITH, B. M. 1990. Mineralogy and geochemistry of epidosites from the Solea graben, Troodos ophiolite, Cyprus. In: MALPAS, J. MOORES, E. M., PANAYIOTOU, A. & XENOPHONTOS, C. (eds) Ophiolites-Oceanic Crustal Analogues. Proceedings of the Symposium Troodos 87. Geological Survey Department, Nicosia, Cyprus, 673 682. - - , SMITH, B. M., VARGA, R. J. & MOORES, E. M. 1987. Geometry, conditions and timing of off-axis hydrothermal metamorphism and ore-deposition

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SAUNDERS, A. D. & NORRV, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications, 42, 313-345. SVERJENSKY, D. A. 1984. Europium redox equilibria in aqueous solution. Earth and Planetary Science Letters, 67, 70-78. VALSAMI,E. & CANN, J. R. 1992. Mobility of rare earth elements in zones of intense hydrothermal alteration in the Pindos ophiolite, Greece. In: PARSON, L. M., MURTON, B. J. & BROWNING, P. (eds) Ophiolites and their Modern Oceanic Analogues. Geological Society, London, Special Publications, 60, 219-232. VAROA, R. J. & MOORES, E. M. 1985. Spreading structure of the Troodos ophiolite, Cyprus. Geology, 13, 846-850. VON DAMM, K. L. 1990. Seafloor hydrothermal activity: Black smoker chemistry and chimneys. Annual Review of Earth and Planetary Science, 18, 173-205. WHITFORD, D. J., KORSCH, M. J., PORRITT, P. M. & CRAVEN, S. J. 1988. Rare-earth element mobility around the volcanogenic polymetallic massive sulfide deposit at Que River, Tasmania, Australia. Chemical Geology, 68, 105-119. WOOD, S. A. 1990. The aqueous geochemistry of the rare-earth elements and yttrium 2. Theoretical predictions of speciation in hydrothermal solutions to 350°C at saturation water pressure. Chemical Geology, 88, 99-125. - ~; WILLIAMS-JONES, A. E. 1994. The aqueous geochemistry of the rare-earth elements and yttrium 4. Monazite solubility and REE mobility in exhalative massive sulfide-depositing environments. Chemical Geology, 115, 47-60.

Drilling of sediment-hosted massive sulphide deposits at the Middle Valley and Escanaba Trough spreading centres: ODP Leg 169 R. H. J A M E S a, R. C. D U C K W O R T H SHIPBOARD

2, M. R. P A L M E R a & T H E O D P L E G 169 SCIENTIFIC PARTY

1 Department o f Earth Sciences, Bristol University, Queen's Road, Bristol BS8 1R J, U K 2 School o f Earth Sciences, James Cook University, Townsville, Queensland 4811, Australia

Abstract: Massive sulphide deposits actively forming from hydrothermal systems within sedimented environments have been drilled during Ocean Drilling Program Leg 169 at two locations along the Juan de Fuca/Gorda spreading centres. The Bent Hill Massive Sulphide and Ore Drilling Program deposits, Middle Valley, include iron- and zinc-rich massive and semi-massive sulphides underlain by a well-developed feeder zone characterized by sulphide impregnations and cross-cutting copper-rich veins. Ridge-parallel normal faulting is probably involved in providing high-permeability pathways for focused discharge at the seafloor, and this is a key element in creating these large ore deposits. In strong contrast, massive sulphide recovered from the Central Hill hydrothermal site, Escanaba Trough, suggests mineralization forms only a thin (5-15m) veneer over the sediment sequence. Interstitial waters recovered from this area have chlorinities both significantly higher and lower than seawater. The only way to explain this variation is that the fluids contain a hydrothermal component which has undergone supercritical phase separation at depth. Diffuse discharge of hydrothermal fluids through the sediments evidently precludes the formation of a large ore deposit in this area.

Ocean Drilling Program (ODP) Leg 169 was the second leg of a two leg programme to investigate the geological, geophysical, geochemical, and biological processes at sediment covered spreading centres in the northeast Pacific ocean. The study of sulphide deposits forming within modern day sedimented-ridge environments is particularly important as many world class basemetal deposits in the ancient geological record are hosted within clastic sedimentary sequences (e.g. Sullivan, Mount Isa, Century). Models for the formation of these modern day deposits have previously relied heavily on systematic geochemical and mineralogical investigations of samples from the mound surface, only inferring subsurface processes (e.g. Fouquet et al. 1993; Zierenberg et al. 1993). It is only within the last few years that this situation has begun to change and details have emerged about the three dimensional nature of these h y d r o t h e r m a l mounds through deep-sea drilling (e.g. ODP Leg 139-the associate leg to Leg 169, Mottl et al. 1994, and ODP Leg 158; Humphris et al. 1995). This paper presents some of the preliminary results from the ODP Leg 169 which drilled the Bent Hill massive sulphide (BHMS) and ore drilling program (ODP) m o u n d at Middle Valley, Juan de Fuca ridge, and massive sulphide deposits at Central Hill, Escanaba

Trough (Fig. 1). Most of the discussion focuses on studies performed by the Scientific Party onboard ship; these results are supplemented by post-cruise studies by the first two authors. Firstly, the lithostratigraphic sequence of the holes drilled at the two sites is detailed. The sulphide mineralization encountered during drilling, is then discussed with an emphasis on the relation of these sulphide deposits to the geological record. Finally, some results of the chemical analysis of pore fluids recovered from sediments which host the sulphide deposits are presented. Very few such pore fluid samples have been collected previously. We use these data to provide a unique insight as to mineral formation within sulphide deposits, and fluid flow pathways. Further preliminary results from Leg 169 are published in Fouquet et al. (1998).

Middle Valley Regional setting Middle Valley is an axial rift valley at the northern extreme of the Juan de Fuca Ridge, near its intersection with the Sovanco transform fault (Fig. 1). Middle Valley is a medium-rate spreading centre (58 mm a l; Davis & Villinger 1992), but its proximity to the cold Explorer

JAMES,R. H., DUCKWORTH,R. C., PALMER,M. R. & THE ODP LE6 169 SHIPBOARDSCIENTIFICPARTY. 1998. Drilling of sediment-hostedmassive sulphidedepositsat the Middle Valley and Escanaba Trough

spreading centres: ODP Leg 169. In. MILLS,R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 177-199

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130

125

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Fig. 1. Location map showing the sediment-covered spreading centres at Middle Valley and Escanaba Trough on the Juan de Fuca-Gorda spreading system.

plate results in reduced magma supply and a slow-spreading ridge morphology with a deep and wide axial trough. During the Middle Brunhes Epoch, less than 20 000 years ago and probably within the last 10000 to 15000a, spreading has shifted westward from Middle Valley to West Valley (Karsten et al. 1986; Davis & Villinger 1992), and current magmatic activity is mostly confined to the West Valley. The axial valley is filled to a depth of more than 2 km in places with turbidite sediment shed from the North American continental margin during the Pleistocene glaciation (Karlin et al. 1992). The youth of the oceanic crust, which is probably no more than 400 000 years old (Davis & Villinger 1992), and the thick sediment cover, combine to produce high basement temperatures and accompanying hydrothermal circulation within the sediment and shallow basement. Middle Valley is highly asymmetrical. The western two-thirds is a deep rift that contains the thickest sediment, especially towards the north, while the eastern third overlies a basement bench that is covered by only a few hundred metres of

sediment. This bench is a fault block that lies between two west-facing normal faults. The basement of this fault block is interpreted to consist of a sediment-sill complex lying some 120 m below seafloor (Rohr & Schmidt 1994). Within the eastern third of Middle Valley there are numerous small, circular hills; Bent Hill is one such mound and is located 3 km west of the eastern rift-bounding normal fault. It is approximately 400 m in diameter and 50 m high. A late Pleistocene to Recent basaltic intrusion was intersected beneath this by drilling during ODP Leg 139 and is interpreted to be responsible for uplift of the sediment above the surrounding turbidite plane (Rohr & Schmidt 1994). A massive sulphide deposit is located on the south flank of Bent Hill and is referred to as the BHMS deposit. This deposit is surficially weathered to iron oxyhydroxides and partially buried by sediment. Extrapolation of the spreading rate suggests that the deposit is located over crust that is approximately 320 000 years old (Davis & Villinger 1992). A second sulphide mound (referred to here as the ODP deposit) occurs approximately 350m south of the BHMS along the trend of the N-S scarp that bounds the western side of Bent Hill. The morphology, degree of oxidation, and sediment cover indicate that this deposit is younger than the BHMS deposit. A single 264°C hydrothermal vent is present on the northern flank (Butterfield et al. 1994). Contoured heat flow data for the Bent Hill area show high values centred around this active vent (Davis & Villinger 1992). The southern part of the deposit has yet to be explored. Drilling operations

The BHMS was first drilled in 1991 on ODP Leg 139 (Davis et al. 1992). Two holes (Holes 856A and 856B) were drilled at the top of Bent Hill and six holes (Holes 856C-856H) were drilled on the BHMS deposit (Fig. 2). During Leg 169, a further six holes (Holes 1035A, C, D, E, F and G) were drilled in the vicinity of the BHMS. Additionally, Hole 856H, which had been drilled on Leg 139 to a depth of 93.8 m before abandonment due to unstable hole conditions (Davis et al. 1992), was deepened to 500 metres below seafloor (mbsf). One hole, Hole 1035H, was sited on the ODP sulphide mound. The locations of these drill holes at Site 1035 are also depicted in Fig. 2. Core recovery varied between 5 and 100%, depending on the nature of the material recovered. Drilling at Site 1035 created two new hydro-

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169

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Lithostratigraphic s u m m a r y Eight lithologic units have been recognized during drilling of the BHMS deposit. These are described below. A more detailed study of the sulphide mineralization is provided in a separate section. Figure 3 denotes a generalized stratigraphic summary for each hole; these have been prepared using shipboard visual core descriptions, augmented by selected shipboard petrographic and X R D studies, and by biostratigraphic and sedimentologic studies. Lithologic units, based on recovered core, are apportioned

over the drilled interval to create these generalized, continuous sections. Turbidites and hemipelagic sediments. Sediments recovered in the Bent Hill area are characterized by Holocene and Pleistocene hemipelagic and turbiditic deposits. Sediment overlying the shallow clastic sulphide (upper 2 - 4 m of Holes 1035D and 1035E) is unaltered silty clay, while sediment lying below the clastic sulphides is characterized by alternating hemipelagic and turbiditic sediments of Pleistocene age (Fig. 4a). These sediments commonly contain hydrotherreal alteration assemblages, including carbonate concretions, authigenic anhydrite, and minor amounts of sulphide mineralization. Sediments from the deeper parts of Holes 856H, 1035A, 1035D, 1035F, 1035G and 1035H, are charac-

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massive and semi-massivesutphides sulphide-veinedsediment (10-50% sulphide) sediment with sulphideveinsand/or impregnations(2-t 0% sulphide) sulphideb ~ i m p r e g n a t e d sediment (10-50% sulphide):Deep Copper Zone (DCZ) basaltic sill/sediment complex basaltic flows no core recovery

wash core

Fig. 3. Stratigraphic sections for Holes 856H and 1035A, C, D, E, F, G, and H showing the distribution of mineralised and nonmineralised subunits. The upper 100m of Hole 856H was drilled during Leg 139; data are from Davis et al. (1992). Core depths are reported as metres below sea level (mbsl).

terized by chloritization and silicification of the sedimentary rocks. Correlation of the sedimentary succession between the holes drilled in the vicinity of the B H M S is difficult. Lateral facies variations occur on a scale of a few tens of metres or even of a few metres, and turbidite beds are generally thin ( < 10 cm) or very thin and have little lateral continuity. The Middle Valley depositional system was probably supplied by turbidite currents that originated by slope failure of the western Canada shelf, north of the Columbia River delta, rather than from a single delta point source (Karlin et al. 1992). The sedimentation in the Middle Valley area is evidently influenced by topographic variations that control the sedimentation at various scales. Formation of transform walls or rift fault escarpments can prevent turbidite currents from reaching the spreading centre, thus lowering the sedimentation rate for substantial amounts of time. Smaller scale morphological variations caused by lava flows, growth of hydrothermal mounds, or the uplift of sediment hills similar to Bent Hill may also influence the deposition of turbiditic currents resulting in a lack of lateral continuity.

Clastic sulphides. The uppermost 8.9 m of Hole

1035A recovered a sequence of unconsolidated, black, fine-grained, sulphide-rich intervals interbedded with grey silty turbidites and green hemipelagic mud (Fig. 4b). This unit is designated clastic sulphides, and is also found between 2.69 and 7.12 mbsf at Hole 1035D, and between 3.37 and 10.12 mbsf at Hole 1035E. Further downcore in Hole 1035D there is a second, older interval of clastic sulphide sedimentation (35.6-36.75 mbsf). These clastic sulphides are interpreted as erosional products shed from the BHMS deposit. The terrigenous nature of the majority of the non-sulphide sediment in the clastic sulphide beds implicates extra basinal-derived turbidity currents as a possible triggering mechanism for accelerating the mass wasting of the sulphide debris. A second type of clastic sulphide was recovered from the top of Holes 1035C, 1035F, and 1035H. These are fragmental sulphides comprising rubble and sulphide breccias, and are interpreted to be formed on the top of the sulphide mound from disaggregated chimneys and reworked m o u n d crusts. The sulphide breccia developed in this way at the top of Hole

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 1035H (the ODP mound) is particularly well developed and is approximately 18m thick, which suggests long-lived venting and sulphide precipitation on the seafloor.

Massive and semi-massive sulphides. Massive sulphide was intercepted in the BHMS deposit in Holes 856H, 1035A, 1035C, 1035D, 1035F, and 1035G during Leg 169, and also encountered in a further 6 holes during Leg 139 (Holes 856C-H). These holes constrain the vertical and lateral extent of sulphide mineralization within the BHMS deposit. Massive sulphide extends at least 40m north, 60m south, and 75m to the west and east of Hole 856H which is located on a topographic high. The vertical thickness of the massive sulphide deposit at 856H is 103.6m. The most common sulphide facies recovered from the BHMS during Leg 169 is massive vuggy and colloform pyrit~marcasite (with or without sphalerite and/or chalcopyrite; Fig. 4c). This facies is considered to be the precursor through which later hydrothermal fluids have flowed and precipitated the iron, copper, and zinc sulphides seen in areas of intense palaeohydrothermal flow near the centre of the deposit. Two other sulphide facies are recognized at the BHMS; a fine grained homogenous massive pyrrhotite, and a heterogeneous pyrrhotite pyrite (with or without magnetite, sphalerite, and copper-iron sulphides). Three separate stacked sequences of massive sulphide underlain by feeder zone mineralization were encountered at the ODP mound in Hole 1035H. The deepest mineralization from this 248 m deep hole was recovered from 162 mbsf. The most common sulphide facies recovered is a sphalerite-pyrrhotite-pyrite-magnetite assemblage comprising up to 50% Zn. This assemblage is far more zinc-rich than any recovered from the BHMS. The sphalerite is either massive and fine-grained or coarse-grained and vuggy, and is commonly paragenetically later than pyrite and pyrrhotite. The most common copper minerals are isocubanite and chalcopyrite. Sulphide feeder zone. Sulphide feeder zones stratigraphically underlie the massive sulphide horizons at the BHMS and ODP deposits. They consist of veins and impregnations of dominantly pyrrhotite and isocubanite that vary in thickness and orientation (Fig. 4d). The feeder zone underlying the BHMS deposit is ~ 1 0 0 m thick, about the same thickness as the massive sulphide deposit itself. It can be divided into three subunits (Fig. 3) defined on the basis of abundance of sulphide minerals and vein morphology. The upper unit ('sulphide veined

181

sediment') comprises 10-15% by volume sulphides, and is dominated by vertical to subvertical crack-seal veins of pyrrhotite and isocubanite that crosscut strongly altered mudstone and siltstone. The middle subunit ('sediment with sulphide veins and/or impregnations'), has significantly fewer veins and generally represents weaker feeder zone mineralization (2-10% by volume sulphide minerals). This unit is characterized by sulphide blebs that have infilled and grown out from pore spaces. The lower unit ('sulphide banded/impregnated sediment'), is distinct from the other two subunits in that it contains 10-50%, by volume, sulphide minerals, mostly as subhorizontal isocubanite-chalcopyrite replacements of original sedimentary structures in a strongly quartz~zhlorite altered sandstone unit. This unit is particularly copper rich, and is referred to as the Deep Copper Zone (DCZ). The mode of formation of the DCZ is discussed in Zierenberg et al. (1998). No veining was observed in this unit. Immediately above this unit there is an intensely silicified zone. The sulphide feeder zone is generally well developed immediately beneath the BHMS except in the westernmost holes (Holes 1035A and 1035G). Correlations of the BHMS feeder zone to the south are unclear because most of the subunits do not correspond directly with the ODP mound feeder zone subunits.

Basaltic sill/sediment complex. Igneous rocks were encountered in Hole 856H at a depth of 432 mbsf. Five basaltic units, separated by finegrained sedimentary rocks, were recognized from the cores recovered, from changes in drill penetration rates, and downhole formation microscanner (FMS) logs. The intercalated igneous bodies are interpreted to be sills, but the possibility that these units could be flows, separated by sediment cannot be discounted. FMS logs indicate the presence of pillow-lava structures in the lowermost unit which may suggest that this erupted onto the seafloor as a flow. The core recovery suggests thickness variation in the sills from a minimum of 0.3 m to a maximum possible thickness of 7.2m. No regular progression of sill thickness with depth is apparent. The sills are typically separated by thin intervals of greenish-grey mudstone and interbedded mudstone and siltstone, and their emplacement has resulted in mineralization along fractures in the uppermost mudstone. The sills are aphyric to highly clinopyroxeneand/or sparsely plagioclase-phyric basalts which are cryptocrystalline (particularly in the thinner

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Fig. 4. Examples of different rock types comprising the sub-surface portion of the BHMS. (a) Interbedded turbidites and hemipelagic sediments, typical of those overlying massive sulphide. This example is from Hole 1035A, 24.5 mbsf. (b) Clastic sulphides (example from the near-surface of Hole 1035A). (e) Massive sulphide: this is an example of the most common type of massive sulphide recovered from the BHMS. It is characterized by massive pyrite and marcasite, and is commonly vuggy (10-50% vugs) with wavy colloform banding on a submillimetre to centimetre scale. This example is from Hole 1035C, N40 mbsf. (d) Example of sulphide-veined sediment which comprises part of the feeder zone underlying the BHMS. The major sulphides present are isocubanite, chalcopyrite, and pyrrhotite, with minor pyrite. This example is from Hole 856H, ~120 mbsf. (e) Variolitic basalt from the base of Hole 856H. The basalt is highly altered, with some fresh plagioclase. (f) Coarse-grained recrystallised massive sulphide from a deep sulphide lens in the ODP deposit. (g) Fluid channelway (vertical) in colloform and vuggy pyrite. Marcasite lines the channelway walls.

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units or near chilled margins) to mediumgrained. Finer-grained margins are present in all five sills, and both the upper and lower chilled margin were commonly recovered. Hydrotherreal alteration of the sills is moderate to complete. The basaltic groundmass is commonly bleached by alteration or tinted green due to chloritization.

Basaltic flows. The deepest sedimentary rock in Hole 856H was recovered from 471 mbsf. Below this depth, c. 29m of volcanic rocks were penetrated. These rocks are moderately to highly altered, glassy to medium-grained, aphyric to sparsely pyroxene- and/or plagioclase-phyric basalts and are macroscopically similar to the igneous rocks that intrude the overlying sediments (Fig. 4e). The presence of altered volcanic glass requires that these rocks were erupted directly on the ocean floor, or intruded in the shallow subsurface and rapidly quenched by seawater. There is insufficient penetration in Hole 856H to establish unequivocally whether these flows erupted directly onto the basaltic ocean crust and hence represent the uppermost basement rocks in Middle Valley, or if there are further sedimentary rocks at deeper levels.

Sulphide mineralization and geochemistry SulphMe facies. Massive sulphides recovered from Holes 856H, 1035A, 1035C, 1035D, 1035F, 1035G, and 1035H can be broadly classified into four facies which are described here. Descriptions of the sulphides recovered on Leg 139 from the main body of the BHMS are provided in Davis et al. (1992) and Duckworth et al. (1994). Fine-gra#wd homogeneous massive pyrrhotite. Massive pyrrhotite is the d o m i n a n t facies recovered from Hole 856H on both Legs 139 and 169.The first core drilled at Hole 856H on Leg 169 was a wash core recovered from 93.8 m which consisted of three pieces of massive pyrrhotite with pyrite and up to 10% magnetite. These pieces are probably rubble that collected at the bottom of the hole during Leg 139 drilling or during Leg 169 re-entry and are similar to massive fine-grained pyrrhotite collected from the b o t t o m of the hole during Leg 139. Subsequent cores recovered massive sulphide to a depth of 103.6m in Hole 856H. The massive pyrrhotite is partly replaced and veined by coarser-grained pyrite and some areas contain pyrite blebs that have formed after the pyrrhotite. The last piece of massive sulphide recovered was a pyrrhotite breccia consisting

of sub-rounded clasts of massive pyrrhotite in a matrix of isocubanite/chalcopyrite with minor pyrite, anhydrite, and amorphous silica. The clasts of pyrrhotite are recrystallized into a granoblastic texture with minor interstitial chalcopyrite and pyrite. Pyrrhotite clast margins are commonly replaced by pyrite and marcasite aggregates. Anhedral, fine-grained magnetite is associated with this pyrite, but does not occur in the unaltered pyrrhotite clasts. The chalcopyrite and isocubanite in the breccia matrix form lamellar intergrowths governed by crystallography of the minerals. Massive pyrrhotite samples were also recovered from the edge of the Bent Hill sulphide mound in Holes 1035C and D (to the south and east of Hole 856H). Additionally, rubble samples in some cores from these peripheral holes consist of siltstone and mudstone fragments that are partially to completely replaced by pyrrhotite.

Heterogeneous pyrrhotite-pyrite + magnetite, sphalerite, Cu-Fe sulphides. This type of sulphide mineralization is similar to many of the samples that were collected on Leg 139 from throughout the BHMS. A common characteristic of this type is the occurrence of fine- to coarse-grained pyrite and pyrrhotite with minor magnetite, sphalerite, and chalcopyrite. Despite being abundant in cores from the centre of the BHMS recovered on Leg 139, on Leg 169 this massive sulphide facies was recovered only from 1035D at ~80 mbsf, and is best described as heterogeneous fine-grained pyrrhotite-pyrite-magnetite with m i n o r ( ~ 1 % ) c h a l c o p y r i t e and anhydrite. However, in Cores 1035D 10X-1 and 1035F 9R-1 sulphide-sediment rocks were recovered that are characterized by subequal amounts of fine-grained pyrite and pyrrhotite that cross-cut and impregnate hydrothermally altered mudstone. The sulphide content generally varies between 50 and 70% by volume and the remaining altered mudstone clasts rarely exceed 5cm in diameter. Pyrrhotite is locally partly replaced by pyrite, and pyrite also rims altered sediment clasts. Magnetite and chalcopyrite are present within the pyrite-pyrrhotite intergrowths and also in millimetre veinlets that cross-cut both the pyrite-pyrrhotite and altered sediment. Pyrite and magnetite occur in a finegrained intergrowth which suggests that they coprecipitated, possibly as a result of the oxidation of primary pyrrhotite. A maximum of 2% sphalerite occurs in this facies.

Sphalerite-pyrrhotite-pyrite-magnetite. Sphalerite-pyrrhotite-pyrite-magnetite massive to semi-

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 massive sulphide is the most common sulphide facies in the ODP mound (Hole 1035H). In contrast, the only zinc-rich interval recognized from the BHMS deposit during Legs 139 and 169 was between 25 and 30 mbsf in Hole 856H, where post-cruise analysis indicated 2-8% zinc in some samples (R. Zierenberg pers. comm. 1996). This Zn-rich mineralization occurs at five different intervals in Hole 1035H, with the main concentrations in the uppermost 26 m (a sulphide breccia unit composed of disaggregated chimneys and reworked mound crusts) and in three intervals between 123 and 162 mbsf (massive and semi-massive sulphides with altered sediment). A thin interval (35cm recovered) of this same type of s p h a l e r i t e mineralization is also present at a depth of 74.7-75.1 mbsf. The three lenses between 123 and 162 mbsf are dominated by black, iron-rich sphalerite and pyrrhotite or pyrite. The sphalerite content is typically between 40 and 60% by volume and ranges up to 70-80% by volume. The iron content of this sphalerite is high, averaging 1112% Fe by weight. These sphalerite-rich parts of the massive sulphide are either compact, earthy and fine-grained, or vuggy and coarse-grained, suggesting two separate precipitation events. In some of the vugs, zinc sulphide is present as hexagonal crystals which could be wurtzite or sphalerite after wurtzite. The predominant copper-sulphide is an intergrowth of chalcopyrite and isocubanite which occurs in minor amounts with subordinate chalcopyrite. Cu-Fe sulphide and locally marcasite occur as finegrained inclusions in the sphalerite-rich areas. Cu-Fe sulphide also forms trails along the crystallographic directions in the sphalerite, and also maps out primary sphalerite/wurtzite grain boundaries. Galena occurs as rare inclusions in the sphalerite. Pyrite is uncommon, and where observed, replaces pyrrhotite. Magnetite is generally associated with pyrite. Pyrrhotiterich parts of this subunit commonly form an open network texture, with interstitial sphalerite, and voids filled with a white and greenish mineral, probably a mixture of talc, smectite and/or chlorite. Some samples have up to 10% interstitial white clay minerals. Chlorite is the most abundant non-sulphide, typically forming a reticulate boxwork pattern interstitial to the sulphides. The only other non-sulphide recognized is carbonate, probably dolomite, as an aggregate in one of the sections. In the intervals of massive to semi-massive sulphide at 134.5-136.4 mbsf, samples are richer in pyrite and magnetite compared to sphalerite and pyrrhotite. Characteristic of this zone is the

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abundance of very coarse-grained neoblastic pyrite (up to 1.5cm diameter), which results in a spotted texture (Fig. 4f). This coarse-grained pyrite has locally been fractured and infilled by later pyrrhotite. Pyrrhotite forms up to 1 mm large skeletal and poikiloblastic grains and aggregates, with inclusions of sphalerite, magnetite and chlorite. This is evidence of fast crystal growth. Similarly, magnetite forms partly euhedral, poikiloblastic, diamond-shaped grains with sulphide inclusions. Magnetite has partially replaced haematite, and earlier formed globular aggregates of haematite laths have locally been totally replaced by magnetite. Sphalerite occurs as subhedral cubes, sometimes with an inclusionfree rim, surrounding Cu-Fe sulphide-diseased core. Carbonate (probably dolomite), quartz and fibrous talc are interstitial to the sulphides. The factors which resulted in the coarse recrystallization of pyrite and the other sulphides in this interval are not known. The presence of epidote locally indicates higher temperatures and/or greater degrees of waterrock interaction, possibly due to incursions of heated, modified seawater into the hydrothermal circulation system or nearby presence of igneous rocks not intersected by this hole. Green hedenbergite (confirmed by XRD) was found in Core 169-1035H-16R-2 at 25 cm and supports this assumption.

Massive vuggy and colloform pyrite-marcasite (±sphalerite chalcopyrite). This is the most common sulphide facies recovered from the BHMS deposit during Leg 169. It is characterized by massive pyrite and marcasite that is most commonly vuggy (10-50% vugs) with wavy colloform banding on a submillimetre to centimetre scale (see Fig. 4c). A mesh network of finegrained pyrite enclosing open space results in a reticulate texture. These vugs vary in size and shape but are commonly irregular and between 0.1 mm to 2cm long. A salient difference between sulphides of this facies recovered during Leg 169 compared with those recovered on Leg 139 is the presence of anhydrite and carbonate which partly fills the vugs. Although this sulphide type is common in Holes 139856 G and H, the vugs in these samples are commonly filled with white amorphous silica or smectite. However, the non-sulphide minerals noted infilling vugs in samples recovered on Leg 169 are clear to grey anhydrite which, in places, occurs as euhedral crystals up to 2 cm long, and coarsegrained carbonate. Vugs are generally lined with euhedral pyrite crystals. Vertical fluid channelway structures are common (Fig. 4g) and channelway walls are commonly lined with

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coarse-grained marcasite. Some samples of massive and vuggy 'pyrite' contain up to 50% marcasite with secondary pyrite. This facies type is also locally enriched in sphalerite and/or chalcopyrite which was not observed in samples from Leg 139. Brown or black sphalerite is present as colloform bands and veinlets that parallel and rhythmically alternate with the pyrite-marcasite bands lining the vertical channelway structures. Bands of dark material (mostly clay-minerals) are also partially intergrown with pyrite. Additionally, black sphalerite locally also occurs interstitial to the pyrite in vugs. Two generations of sphalerite are present; an earlier phase which is coeval with a generation of pyrite that replaces marcasite and which displays extensive chalcopyrite disease (Barton & Bethke 1987) and later, less Fe-rich sphalerite, that occurs in veinlets cross-cutting pyrite. Pyrrhotite and isocubanite also occur in these latter veinlets. In Hole 1035C, massive vuggy pyrite was recovered that contains up to 20% chalcopyrite as vug infill and cross-cutting veins. It is likely that this massive marcasite-pyrite facies is the precursor through which later hydrothermal fluids have flowed and precipitated later iron, copper and zinc sulphides. In areas of intense palaeohydrothermal fluid flow, this facies has been mineralogically modified to produce the distinctly different sulphide types recovered in Holes 856H and 1035H. Bulk metal distribution in sulphides. A series of massive and semi-massive sulphide, clastic sulphide, and feeder zone sulphide samples from the BHMS and ODP mound were analysed onboard ship for their metal (Cu, Zn, Pb, and Fe) and sulphur contents in order to obtain a preliminary picture of the range and maxima of metal grade, and of metal ratios for comparison with ancient economic mineral deposits. Analytical methods are described in Fouquet et al. (1998). These spot analyses cannot be used to calculate true metal grades, which requires continuous sampling of intervals, but the samples were chosen to be representative of different styles of mineralization and to test visual estimates of Cu and Zn abundances. Sulphides recovered from the BHMS deposit are dominated by pyrite and pyrrhotite and base metal concentrations are relatively low. Highest concentrations of Zn (up to 6.1% by weight) are found in the massive sulphides, while Cu concentrations are highest (up to 1.6% by weight) in the feeder zone, especially at its base in the sulphide-banded sediments of the DCZ. Pb averages 180 ppm in the BHMS, and is

highest (up to 730 ppm) in clastic sulphides from the mound subsurface. Compared with ancient massive sulphide deposits, the BHMS is a Pbpoor, Fe-rich, Cu-Zn deposit with metal ratios similar to many Besshi-type deposits (Slack 1995). Massive sulphide recovered from the ODP mound is diverse and highly heterogeneous. For the eight samples analysed, the average Zn is 19.6% by weight and the average Cu is 3.3% by weight, but the variance is large. Zn in massive sulphide ranges up to 51% by weight, and Cu, which is highest in the feeder zone, ranges up to 16.6% by weight. Pb in the ODP mound averages 132 ppm. These data indicate that there is zoning of metals within the BHMS deposit. Cu is significantly enriched in the feeder zone, Zn shows highest concentrations in the massive sulphide deposit and Pb shows some enrichment in the uppermost clastic sulphides relative to the rest of the deposit. These zoning patterns agree well with the known solubility behaviour of Cu, Zn, and Pb in chloride solutions with respect to sulphide formation along a cooling trend (Johnson et al. 1992; Seyfried et al. 1991; Shanks & Bischoff 1977). However, there is abundant textural evidence of complex and heterogeneous hydrothermal reworking within the deposit (Duckworth et al. 1994) and precipitation mechanisms and flow patterns are not simple, but the basic Cu, Zn, and Pb chloride complexing controls on metal precipitation have probably influenced the final metal distribution patterns. In the complex ODP mound, the deep zinc-rich massive sulphide zone overlies the copper-rich feeder zone, however, the sphalerite commonly replaces primary iron sulphides suggesting multiple episodes of hydrothermal sulphide precipitation. Geochemhstry o f interstitial fluids

A series of interstitial water samples was collected from Site 1035. The chemical composition of these fluids is similar to seawater, and displays only minor changes with depth which are largely related to the early diagenesis of organic matter. However, the lowermost sample collected from Hole 1035A (168 mbsf), and the sole sample collected from Hole 1035H (55 mbsf) have a notably different chemical composition. For example, Fig. 5 (Hole 1035A) shows a steep decrease in C1 from its seawater value (555 mM) to 440 mM below ~150 mbsf, which is accompanied by a steep increase in Li from its seawater value (26#M) to 290#M. These concentrations are similar to those measured in

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 CI (mM) o

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Fig. 5. Pore fluid (a) C1 and (b) Li, Hole 1035A, Bent Hill, Middle Valley. Data for hydrothermal fluids are from Butterfield et al. (1994). Note the deepest sample recovered from this Hole has a very different composition from overlying samples. This sample is isolated by a layer of intensely silicified turbiditic mudstone which acts as an impermeable barrier to upward fluid flow. actively venting hydrothermal fluids collected by submersible in 1990 from Bent Hill: 428 mM C1 and 380 #M Li (Butterfield et al. 1994). Ca, Na, K, B, NH4, and the Na/C1 ratio also exhibit marked changes towards values measured in BHMS hydrothermal fluids in this lowermost sample. What is significant is that a thin layer of intensely silicified turbiditic mudstone was en-

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countered overlying this sample at around 160 mbsf. The pore fluid data suggest that this acts as an impermeable barrier to upward fluid flow, hence the virtual absence of an influence of the vent fluid on the concentration profiles above this horizon. Lateral flow of BHMS hydrothermal fluid can best explain the sudden change in fluid composition. Similarly, the only sample collected from Hole 1035H (ODP mound) has a fluid composition consistent with the addition of a Bent Hill-type hydrothermal fluid (Fig. 6). Of importance is that drilling at this site re-initiated hydrothermal venting, suggesting that, although no silicified horizon was encountered in this hole, drilling again penetrated some barrier that prevents fluid from the hydrothermal reservoir from reaching the seafloor (Davis & Fisher 1994). Similar plots are obtained for Ca, Na, K, B, NH4, and the Na/ C1 ratio. These results suggest that both the BHMS and ODP mounds are currently underlain by identical hydrothermal fluids, which has implications for the lateral continuity of the deposits. However, several lines of evidence suggest that the BHMS deposit formed at least 10000 years ago (Davis et al. 1992), hence the hydrothermal fluids presently underlying the deposit cannot be part of the same system that precipitated the BHMS deposit. In this connection, it is surprising that there is an abundance of anhydrite in the vuggy pyrite around the edges of the BHMS in Holes 1035A, 1035C, 1035D, and 1035F, as anhydrite dissolves below 150°C, and is thus unstable. This suggests that this mineral phase may be much younger than the BHMS itself. As discussed previously, anhydrite was not found as an interstitial phase in samples from Leg 139 (in 1991) from the centre of the BHMS deposit. Thus we believe that there may be a new hydrothermal regime in the Bent Hill area which precipitates anhydrite in the ancient sulphides at the BHMS mound and is associated with the ODP mound circulation system. Summary

The holes drilled through the BHMS on Leg 169 and Leg 139 constrain the minimum extent of the BHMS to at least 100 m in the N-S direction and 150 m in the E-W direction. The thickness at the centre of the mound is 103 m. There is an abrupt contact underlying the massive sulphide with a sulphide feeder system which is up to 107m thick in Hole 856H and represents the pathway for the hydrothermal fluids that form the BHMS deposit. The feeder system is welldeveloped indicating that focusing of hydro-

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thermal discharge is a key element in creating the Middle Valley ore deposit. In this connection, FMS logs of Hole 856H identified two zones (221-239 and 250-270 mbsf) with low resistivity and very high porosity which are characteristic of fault zones or fragmented formations (Fouquet et al. 1998). The top of these two intervals dips ~50°W, strongly suggesting that ridge-parallel normal faulting provides high-permeability pathways for focused discharge at the seafloor. With the exception of a 5 m thick interval near the top of the mound (present in all holes except Hole 1035G), the massive sulphide assemblage contains no intermixed sedimentary component. This isolated interval of sulphides is composed of predominantly sand-to-clay sized sulphide grains with altered blue-green clay similar to that observed in the weathered residue formed by the dissolution of collapsed anhydrite-rich hydrothermal chimneys from the active vent fields (Turner et al. 1993). Thus, this interval probably represents a turbidite that was deposited as the hydrothermal system was active, incorporating sulphide debris shed from the mound including disaggregated chimneys and reworked sulphide crusts. The general lack of interbedded sediment indicates that the bulk of the massive sulphide deposit at Bent Hill formed in a single episode of sulphide mound building that was rapid relative to the rate of turbidite sedimentation. This is discussed further by Zierenberg et al. (1998). In contrast, the lithostratigraphic sequence underlying the active ODP deposit is considerably more complex than the BHMS. The occurrence of stratigraphically stacked massive

sulphide lenses indicates that hydrothermal venting can be episodic at a single discharge site. Flow of hydrothermal fluid through previously deposited sulphide mineralization causes recrystallization and local redistribution of metals which enhances the separation of Cuand Zn-rich mineralization leading to extremely high base metal grades, with some samples exceeding 50% Zn and 16% Cu. The chemical analysis of pore fluids recovered from sediments which host the BHMS and ODP deposits suggests that the present day hydrothermal regime in this part of the Middle Valley is controlled by an impermeable silicification horizon at depth. The composition of pore fluids from immediately below this horizon closely matches the composition of hydrothermal fluids collected from high-temperature vents on the flank of the ODP mound. This suggests that the hydrothermal fluids have not reacted extensively with the sediments and they are strongly focused in zones of high permeability.

Escanaba Trough Regional setting

The Gorda Ridge spreading centre is located off the coast of Oregon and northern California and is bounded by the Mendocino Fracture Zone on the south and the Blanco Fracture Zone on the north. A small offset in the spreading axis at 41°40'N marks the northern boundary of the Escanaba Trough, which forms the southernmost part of the Gorda Ridge. The spreading rate is 2 4 m m a -1 (Atwater & Mudie 1973) and the trough has a typical slow-spreading ridge

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 morphology, with an axial valley lying at > 3250 m which is flanked by steep walls which rise 900 to 1500 m above the valley floor. At its northern end, the axial valley is only a few kilometres wide, but it widens to about 18 km at the southern end. The Escanaba Trough is divided into two 80 to 100km long segments separated by a 5km right-lateral offset at 41°08'N. South of 41°17'N, most of the valley is covered by interbedded hemipelagic and turbiditic sediments (Vallier et al. 1973; Clague & Holmes 1987; Normark et al. 1994). At the southern end, the sediment thickness exceeds 900 m. The sediment accumulation rate during the late Pleistocene was as high as 8 m per 1000 a (Normark et al. 1994). During the Holocene, the sedimentation rate averaged 14cm per 1000a (Karlin & Zierenberg 1994). Therefore most of the sediment fill was probably accumulated in less than 200 000 years, during the last two or three sea-level low stands during the Pleistocene. These sediments have sealed off the hydrotherreal circulation that normally cools the oceanic crust. Volcanic and intrusive emplacement of ridgeaxis basalt affects the sediment fill at several localities along the trough (Morton et al. 1987, 1994). The volcanic edifices are characterized by an abrupt transition from undisturbed sediments to deformed, tilted, and commonly uplifted sediment layers (Morton et al. 1994). Associated with the larger scale igneous centres are small circular sediment hills that are ringed by faults and uplifted 50 to IL00m above the surrounding area.

Central Hill is one of these small uplifted sediment hills. It is 1 km in diameter and 60 m high, and is steep sided, formed of uplifted and tilted sediment. Central Hill is interpreted as a cylindrical block of sediment bordered by curvilinear normal faults, and was probably uplifted by the intrusion of a large basaltic sill. The western, sediment-covered part of the Central Hill contains the most extensive sulphide deposits observed in the Escanaba Trough. The massive sulphide deposits on the northwest flank of the Central Hill are actively venting hightemperature (108-217°C) hydrothermal fluids (Campbell et al. 1994). The chemical composition of these fluids is substantially different from sediment-starved mid-ocean ridge (MOR) hot springs due to their reaction with overlying sediments. They are enriched in K, Sr, Rb, Cs, B, and T1; depleted in sulphide-forming metals; have lighter B isotope ratios, and higher radiogenic 87Sr/86Sr ratios than MOR vent fluids (Campbell et al. 1994). A second area on the

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southeast side of Central Hill has abundant clams, and bacterial mats, and issues diffuse, presumably low temperature, fluid from a small mound. No fluid samples have been obtained from this area. Drilling operations

A transect of nine holes (Holes 1038A to 1038I; Fig. 7) was made in the vicinity of Central Hill. Holes 1038A, 1038B, 1038E and 1038H are located on the northwest flank of the hill, close to the area of high-temperature discharge. Holes 1038C and 1038D are on the north flank of the hill. Holes 1038F and 1038G are on the southeast flank of the hill near the area of diffuse flow. Hole 1038I was drilled at the top of Central Hill, between the areas of high temperature and diffuse venting. Core recoveries were low (2.3-17.5%) in the exploratory phase of drilling (Holes 1038A to 1038F), where a rotary coring system was used to facilitate recovery of massive sulphide. In order to improve recovery at depth, the rotary coring system was exchanged for the standard advanced piston core/extended core barrel (APC/XCB) system to drill holes 1038G to 1038I, and core recovery subsequently improved (16-47%). Lithostratigraphic summary

The stratigraphic units intersected in each of the holes drilled at Central Hill are given in Fig. 8. Sulphides were only recovered in the uppermost few metres of cores 1038A, 1038B, 1038C, 1038E and 1038G (Fig. 9a). The mineralogy of these sulphides is discussed in a later section. The remainder of the cores consists of turbiditic and hemipelagic sediments which are highly altered at all sites. Stratigraphic correlation between the holes is therefore extremely difficult. The sediments have tentatively been assigned to 7 lithologic units in Fig. 8 (Units II-VIII). These units are discussed briefly below. Unit II is characterized by graded beds of fine sand to silt and silty clay interbedded with hemipelagic mudstone. At the base of this unit there is an abrupt contact with Unit III, which is characterized by a change to thicker, more sandrich turbidites separated by relatively thin mudstone intervals. Units IV and V are characterized by beds of fine to medium sand, interbedded with siltstone and claystone. A thin basaltic layer was recovered in the Unit IV/V interval in Holes 1038G, 1038H and 1038I; this is discussed in a following section. Unit VI was intercepted only at Hole 1038I, and recovery

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characterized by stacked siltstone turbidites with no obvious intervening hemipelagic intervals. The widespread alteration of sediments observed at Central Hill is of special interest as it pertains to diagenesis, hydrothermal circulation, basalt intrusion and sulphide mineralization. For these purposes, Fig. 8 assigns sedimentary intervals to alteration facies, which are defined solely on the basis of visual core description and smear slide analysis. These assignments are therefore preliminary, and require validation by further chemical and mineralogical studies. Facies 'a' is defined as relatively unaltered primary sediments. Facies 'b' contains authigenic carbonate in the form of cement or c a r b o n a t e nodules. The presence of large amounts of authigenic carbonate in the sediments makes this a useful, though not infallible, indicator of thermally-accelerated rates of diagenesis. Facies 'c' is defined as clay- and chloritealtered, non-calcareous sediments. This alteration facies is generally limited to shallow sediments underlying massive sulphide deposits in holes that are very near to active hydrothermal vents. In Holes 1038A, 1038E and 1038H, the hydrothermal alteration 'c' facies is underlain by the carbonate 'b' facies, a succession that is believed to represent more altered sediments overlying less altered ones. This sequence could result from lateral flow of hydrothermal fluids above a shallow seawater recharge zone. Facies 'd' is used to describe sediments from which all carbonate minerals

41000 '

Fig. 7. Location of holes, active vents, and exposed massive sulphide at Site 1038, Central Hill, Escanaba Trough.

over this interval was extremely poor. The few & strata are dominated by massive, silty claystone. Unit VII is dominated by moderately indurated, calcareous silty claystones. These claystones are interbedded with rare muddy turbidites with thin siltstone bases. Unit VIII is situ

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Fig. 8. Lithologic units and sediment alteration facies recovered from Holes 1038A-I at Central Hill, Escanaba Trough. Lithologic Unit designations and sediment alteration facies are explained in the text.

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 have been dissolved, presumably as a result of replacement or dissolution by advecting hydrothermal fluids. Sulphide mineralization. Sulphides were recovered in Holes 1038A, 1038B, 1038C, 1038E, 1038G and 1038H and comprise mainly fragments from drilling; no oriented pieces were obtained (Fig. 9a). In situ sulphides were recovered from the top of Hole 1038A and 1038E. In other holes, the massive sulphide material is undoubtedly rubble which has fallen in from the top of the holes. Despite this lack of stratigraphic continuity, the fragments obviously represent the massive sulphide that is present at or near the seafloor, and are a useful snapshot as to the mineralogy and textures of the Central Hill sulphide mineralization. The massive sulphide fragments recovered are mainly composed of pyrrhotite with minor pyrite. This is in strong contrast to sulphide recovered from previous cruises in the Central Hill area which recovered samples with a more exotic, polymetallic mineral assemblage with Pb, As and Bi sulphide phases (Koski et al. 1988; Zierenberg et al. 1993). The most indurated massive sulphide recovered on Leg 169 is composed of fine-grained to medium-grained hexagonal pyrrhotite, typically forming an interlocking open network. Variations in the pyrrhotite grain size in a single 1-2 cm fragment is common. Pyrrhotite is locally partly replaced by pyrite, which also occurs as cross-cutting veins and veinlets. Deep red to black sphalerite is abundant in several of the fall-in fragments and occurs mainly intergrown with pyrrhotite, or as interstitial botryoidal masses within an open pyrrhotite network. Marcasite is present in some fragments, associated with pyrite. Minor Cu-Fe sulphide, mainly isocubanite, which can be seen by its colour and rapid oxidation, occurs in some clasts. Several samples have orange goethite or limonite and dark red haematite coatings, often developed preferentially on one side, confirming these samples were weathered at the surface of the mounds, prior to recovery. Native sulphur was identified visually as a weathering product of pyrrhotite. Major gangue minerals are anhydrite and a greenish grey, soft, very fine-grained mineral, probably talc or Mg-smectite. Anhydrite was found only in the areas of active venting. Barite is in places abundant, typically associated with pyrrhotite and sphalerite. Quartz/amorphous silica occurs as a major gangue mineral in a few fragments. The in situ massive sulphides from Holes 1038A and 1038E, and some of the fall-in clasts

191

in the upper cores of other holes, appear to comprise chimney fragments, sometimes with fossil channelways or conduits preserved, and disaggregated mound material. Some chimneytype clasts consist of coarse-grained vuggy pyrite with 2-5 mm conduits lined with finer-grained euhedral pyrite. Vuggy pyrite is not typical for the sulphide deposits at Central Hill that were recorded during previous sampling by the submersibles or dredging. In other pieces, the fluid channelways are defined by variations in the grain size of pyrrhotite. In Hole 1038H two clasts 2-3 cm wide consist of fibrous, columnar aggregates of pyrrhotite which has grown perpendicular to a hydrothermal conduit. Disseminated sulphide in altered sediments was recovered in Holes 1038B, 1038C, 1038E and 1038G. The disseminated sulphides in Hole 1038B are drill cuttings that have fallen in from higher up in the hole, whereas the other occurrences seem to be in situ. The in situ occurrences are mainly weak impregnations (210%) of pyrrhotite with minor pyrite in sand, silt and clay. Minor Cu-Fe sulphide was noted in Hole 1038C. The disseminated sulphides in Hole 1038C are associated with chlorite altered silty clay, whereas no alteration was observed in the 1038G occurrences. A fragment that may have occurred in situ in Hole 1038E (Section 169-1038E-4R-1, 0-8cm) consists of a highly silicified sandstone impregnated by pyrite, marcasite, and possibly pyrrhotite. Pyrite infills vugs and forms a thin crust on the outer edge of the fragment, as well as impregnating the sandstone. This piece immediately underlies the massive sulphide, and is probably part of a thin, poorly developed feeder zone below the deposit. Also, the fragments of very fine-grained chlorite with euhedral pyrite, cross cut by veinlets of anhydrite and minor barite in Hole 1038B, may also represent intense alteration within a feeder zone. Igneous petrology and geochemistry. Basaltic rocks were recovered in Holes 1038H (~135 mbsf), 1038I (~161 mbsf and ~403 mbsf) and 1038G (,-~137 mbsf; Fig. 8). The thickness of these basalt intervals is not well constrained, but based on drilling conditions, the hard basalt layer from Hole 1038H is not more than 1 to 2m thick, and the uppermost basalt layer in Hole 1038I is ~1.5 m thick. Figure 8 demonstrates that the uppermost basalt interval recovered from Holes 1038G, 1038H and 1038I occurs at an approximately similar depth below the seafloor. These upper basalt intersections are thought to be part of a single thin (< 10 m) sill. The basalts are micro-

192

R. H. JAMES ET AL. :::: ::::::.

;

:i;

::

Fig. 9. Examples of different rock types comprising the sub-surface portion of Central Hill. (a) In situ massive sulphide from the surface of Hole 1038A. (b) Microcrystalline, aphyric to sparsely plagioclase-phyric basalt recovered from 135 mbsf in Hole 1038H. The vesicles are lined with chlorite and partly to completely infilled with intergrown white calcite. Numerous narrow (0.2-2 mm) veinlets of chlorite, calcite, and pyrite, cut the basalt. This sample also exhibits a bleached and chloritized chilled margin with adjacent basalt.

crystalline to fine to medium-grained, dark bluegrey, unaltered, moderately plagioclase-phyric and sparsely olivine-phyric (Fig. 9b). Numerous narrow (0.2-2mm) veinlets of chlorite, calcite, and pyrite cut the basalt fragments recovered from this hole. The basalt intersection in the base of Hole 1038I recovered 67 cm of very fine-grained, nonvesicular to sparsely vesicular, sparsely to moderately plagioclase-olivine-phyric and plagioclase-phyric basalt. The upper contact contained a 2 m m thick b a k e d and bleached sediment margin, and the basalt has a ~ 5 m m

thick margin of fresh glass. Based on the presence of this fresh glass and the baked contact, this basalt intersection is thought to be a flow. There is insufficient core penetration from this interval to ascertain if this is igneous basement.

Geochemistry o f interstitial )quids and sediments A suite of pore fluid samples were collected from all holes drilled in the vicinity of Central Hill.

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 The rotary coring system, employed to drill Holes 1038A-1038F, recovered only a limited number of samples suitable for squeezing, but holes drilled using the APC/XCB system (Holes 1038G-1038I) recovered suitable samples at short intervals throughout the core. The following section discusses the results of chemical analysis of pore fluids and their conjugate sediment in cores recovered from Hole 1038I; this hole is the best sampled of the eight, and is representative of all the holes. Figure 10 shows that the chemical composition of pore fluids from Hole 1038I varies significantly from seawater. Fluids collected from a sandy interval (Unit III) lying between 75 and 142 mbsf have concentrations of K, C1, and Na (and Mg and the Na/C1 ratio) which are lower than seawater, while concentrations of Li, (and B, and NH4) are higher. The most striking feature of these data is the wide variation in C1 concentration (300800raM, considering all holes) compared to seawater (555mM). Seawater is the primary reacting fluid in both pore waters and hydrothermal systems, and the concentration of C1 in both basalts and sediments is negligible. This wide variation in chloride is therefore inconsistent with thermal alteration or dehydration and transformation of hydrous mineral phases, or interaction with sill complexes (e.g. Gieskes et al. 1982; Kastner et al. 1991). The only way to explain this variation in C1 is that the pore fluids contain a hydrothermal component which has undergone phase separation in the super-critical region. This leads to the formation of a Cl-poor vapour phase and a Cl-rich brine phase (Bischoff & Rosenbauer 1988). The phase-separated components subsequently re-mix in variable proportions with background pore fluids. As C1 is the principal anion in vent fluids, charge-balance considerations dictate that changes in dissolved C1 can be expected to induce systematic changes in the concentration of dissolved cations. Values of Ca and Na in the pore fluids are clearly strongly determined by C1 (Fig. 11). Low-C1 fluids have low Na and Ca, while high-C1 fluids have high Na and Ca. The relationship between C1 and the minor cations which have high concentrations in sediment relative to basalt is more complicated (Fig. 12). Hydrothermally modified pore fluids from Site 1038 usually have higher concentrations of Li, NH4, and B than vent fluids from both sediment-free and sediment-hosted hydrothermal systems. This suggests that the pore fluids have interacted extensively with relatively fresh sediment at elevated temperatures and a low water/sediment ratio. The concentration of

193

K in the pore fluids is generally lower than in vent fluids from both sediment free and sediment hosted hydrothermal systems, which probably reflects uptake of K during silicification reactions in the sediments associated with the formation of K-feldspar (Kastner & Gieskes 1976). It is interesting to note that the conjugate C1rich and Cl-poor fluids appear to be physically segregated. Low-C1 fluids tend to be concentrated in sand-rich sedimentary facies, and highC1 fluids are focused close to basalt intrusions. This is consistent with a model proposed by Fox (1990) for fluid phase segregation based on relative permeability. In this model, hot fluids from the subsurface migrate under pressure through the most permeable (i.e. sand-rich) horizons. However, the presence of a small percentage of vapour phase effectively reduces the permeability to the brine phase which means that the brine phase tends to be confined to the major conduits, such as a basalt intrusion. The Li, Na, and K content of the conjugate sediment phase has been determined for a small number of samples from Hole 1038I (Fig. 13). The most notable feature of these profiles is a reduction in Li and K, and an increase in Na, in mineral phases recovered adjacent to the sill. These changes are accompanied by large increases in pore fluid Li, K, and Na. This suggests that reactions with sediments are strongly controlling pore fluid levels of K and Li close to the sill, but clearly some other process (i.e. phase separation, as discussed previously) is controlling pore fluid concentrations of Na. Of importance is that we do not see nearly complete depletion in Li in the mineral phases as has been previously noted at Escanaba Trough (Campbell et al. 1994) and in hydrothermally altered sediments from the Guaymas Basin (DSDP Site 477/477A; Chan et al. 1994). This means that uptake of Li into secondary mineral phases is probably important. The increase in Na in the solid phase may be the result of plagioclase recrystallization; Na is fixed from solution during albitization of plagioclase feldspar (Thornton & Seyfried 1987). Summary

The massive sulphides recovered from Site 1038 consist predominantly of massive pyrrhotite or vuggy pyrite. Sphalerite is locally abundant and is associated with pyrrhotite. The poor recovery of sulphides on this leg did not confirm the presence of high Pb, Bi and As contents in the deposits at Central Hill which have been recorded from previous sampling of the area

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Fig. 11. Concentration in pore fluids (from all Holes at Site 1038, normalized to Mg = 0) plotted v. C1. (a) Ca. (b) Na. Endmember concentrations in vent fluids from bare rock and sediment-hosted hydrothermal sites are provided for comparison (data from Von Damm 1990; Campbell et al. 1994; Charlou et al. 1996). Non-zero Mg in the pore fluids is assumed to indicate contamination with seawater during the sampling procedure; temperatures in the sediments from which the pore fluids were extracted generally exceed 160°C (based on analysis of organic components; B. Simoneit, pers. comm.), and at these temperatures Mg is quantitatively removed from circulating fluids (Thornton & Seyfried 1987). As all samples shown in this figure contain < 7 mM Mg, extrapolation results in only small changes to the values.

(Koski e t al. 1988; Zierenberg e t al. 1993). However, the presence of these more exotic phases suggests that the fluids precipitating the Central Hill sulphides have reacted extensively with sediments and have leached metals during this interaction. This is in contrast to the Middle Valley sulphides recovered from Sites 169-856H and 1035H which have a dominant F e - Z n - C u geochemistry, consistent with derivation of metals from a predominantly basaltic source. The paucity of massive sulphide recovered from this site suggests that the mineralization forms only a thin veneer (5-15m) over the sedimentary sequence of Central Hill. No major intersections of massive sulphide were recovered. The recovery of sediments at shallow depth beneath the surface and the absence of a welldeveloped feeder system, rules out the possibility of a thick (more than 20-25 m) sulphide lens in the areas which were drilled. In addition, the absence of a well-developed feeder zone in the sediment under the sulphide mounds suggests a lack of distinct channelways of high permeability (such as a fault zone). This agrees with the lack of chimneys and the general occurrence of high temperature pyrrhotite crusts on the sediment as observed during submersible dives (Zierenberg & Shanks 1994). The chemical analysis of interstitial waters supports the observations described above. Levels of C1 are significantly different from seawater in many samples, indicating widespread contribution of a hydrothermal component. This is consistent with relatively pervasive

circulation and diffuse venting of hot fluid, rather than focused, high-temperature discharge. As the vent fluids mix with background pore waters and cool, a variety of hydrothermal products, including metal sulphides, silica and anhydrite, will be deposited below the sedimentseawater interface preventing a large accumulation of sulphides on the seafloor (Magenheim & Gieskes 1994; James & Elderfield 1996). Pore fluid concentrations of Li, B, and NH4 are far higher than measured in local vent fluids and background pore fluids. This suggests that there is intense thermal alteration of the sediments, with leaching of these elements along with others which may include the more exotic metals (Pb, Bi, and As).

Conclusions A comparison of the results of drilling of the BHMS and ODP deposits, Middle Valley, and the Central Hill sulphide deposit, Escanaba Trough, reveals very different styles of sulphide mineralization. The BHMS and ODP deposits are large sulphide accumulations that are the result of a focused plumbing system with longlived hydrothermal activity. The Central Hill mineralization is a thin veneer at the seawatersediment interface which has formed from pervasive circulation and diffuse venting of hot fluid over a short period of time. This variation in plumbing is strongly controlled by the sedimentary stratigraphy, and it is interesting that within the Central Hill area, the sedimen-

196

R. H. JAMES E T A L . !

7000

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+

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2000

+

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Bare rock l Sediment-hosted Escanaba pore fluids

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50

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200

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600

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Fig. 12. Concentration in pore fluids (normalized to Mg = 0) plotted v. C1. (a) Li. (b) NH4 (e) B (note the logarithmic scale on the y-axis). (d) K. Endmember concentrations in vent fluids from bare rock and sedimenthosted hydrothermal sites are provided for comparison (data source as for Fig. 11). Of importance is t h a t concentrations of Li in the Escanaba Trough pore fluids are very much higher than the general trend depicted in the figure. Levels of Li in pore fluids from the Bent Hill area are very similar to the general trend (see Fig. 5). Note that actual concentrations of B in the pore fluids may be higher than shown as data have not been corrected for temperature of squeezing artefacts (You et al. 1996). tary sequence is m o r e s a n d - d o m i n a t e d t h a n a r o u n d the Bent Hill area. The stratigraphy of the Bent Hill area is such that h y d r o t h e r m a l fluids are focused up fault planes a n d along limited sand-rich facies, thus limiting fluidsediment interaction and forming large sulphide deposits. At Central Hill however, the predominance of permeable sand horizons allows ubiquitous fluid flow and fluid-sediment interaction but with only thin accumulations of sulphides. Chemical analyses of pore fluids from within the sediments which host these deposits support these observations. H y d r o t h e r m a l fluids are f o u n d only below i m p e r m e a b l e horizons at the B H M S and O D P deposits, and do not a p p e a r to have reacted extensively with the sediment; for example, Li concentrations are similar to those recorded in high-temperature fluids issuing from the surface of the O D P m o u n d . In contrast, porefluid chloride concentrations indicate that h y d r o t h e r m a l fluids are ubiquitous in sediments

recovered from the Escanaba Trough. However, this vent fluid signature has been overprinted by thermally enhanced fluid-sediment reactions for those elements, such as Li, which have high concentrations in the sediments relative to basalt. We are grateful to the master, officers, crew and scientific technical support team of the Joides Resolution. This work was supported by NERC through grant BRIDGE 76 and the UK ODP Programme (RHJ), and by an ARC Small Grant (RCD). We thank two anonymous reviewers for their comments.

ODP Leg 169 Shipboard Scientific Party R. A. Zierenberg, Department of Geology, University of California, Davis, Davis, California 95616, USA; Y. Fouquet, IFREMER, BP 70 29280 Plouzan6 Cedex, France; J. Miller, ODP, 1000 Discovery Drive, College Station, Texas 77845-9547, USA; J. M. Bahr, Department of Geology and Geophysics, University of

DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 Li (ppm) 0

20

30

40

50

60

70

100

e-~ 200

g e-

e~

I~1

300

400 Hole 10381

500

(a)

Na (Wt%)

o[ 1

1.25

1.5

,

1.75

2

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~'300

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500 [Hole 1

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K (Wt%) 0 0

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2

3

t

ATWATER, T. & MUOIE, J. D. 1973. Detailed nearbottom geophysical study of the Gorda Ridge. Journal of Geophysical Research, 78, 8665-8683. BARTON,P. B. & BETHKE,P. 1987. Chalcopyrite disease in sphalerite: pathology and epidemiology. American Mineralogist, 72, 451M67.

.~ 2 0 0 -

E

..E: ..i-, ¢)..

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4O0

500

Wisconsin- Madison, 1215 West Dayton Street, Madison, Wisconsin 53706, USA; P. A. Baker, Department of Geology, Duke University, Box 90227, Durham, N. Carolina 27708-0227, USA; T. Bjerkarden, Norges Geologiske Undersokelse, Postboks 3006 Lade, 7002 Trondheim, Norway; C. A. Brunner, Institute of Marine Sciences, University of Southern Mississippi, Stennis Space Centre, Mississippi 39529, USA; R. Gable, BRGM, 45060 Orleans Cedex 2, France; J. Gieskes, Scripps Institution of Oceanography, University of California, San Diego, 9500 Gilman Drive, La Jolla, California 92093-0215, USA; W. D. Goodfellow, Geological Survey of Canada, 601 Booth St., Ottawa, Ontario K1A 0E8, Canada; H. M. Groschel-Becker, RSMAS, University of Miami, 4600 Rickenbacker Causeway, Miami, Florida 33149-1098, USA; G. Guerin, Lamont-Doherty Earth Observatory, Columbia University, Route 9W, Palisades, New York 10964, USA; J. Ishibashi, Laboratory for Earthquake Chemistry, University of Tokyo, 7-3-1 Hongo, Tokyo 113, Japan; G. Itturrino, Lamont-Doherty Earth Observatory, Columbia University, Route 9W, Palisades, New York 10964, USA; K. S. Lackschewitz, Geologisch-Palaontologisches Institut, University of Kiel, Olshausenstrasse 40, D24118 Kiel, Germany; L. L. Marquez, Department of Geological Sciences, Northwestern University, 1847 Sheridan Road, Evanston, Illinois 60208, USA; P. Nehlig, BRGM, BP 6009, 45060 Orleans Cedex, France; J. M. Peter, Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario KIA 0E8, Canada; C. A. Rigsby, Department of Geology, East Carolina University, Greenville, N. Carolina 27858, USA; P. Schultheiss, GEOTEK Ltd., 3 Faraday Close, Daventry, Northants N N l l 5RD, UK; W. C. Shanks, USGS, Box 25046, MS 973, Denver, Colorado 80225, USA; B. R. T. Simoneit, College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon 97331-5503, USA; M. Summit, School of Oceanography, Box 357940, University of Washington, Seattle, Washington 98195, USA; D. A. H. Teagle, Geological Sciences, 2534 C. C. Little Building, University of Michigan, Ann Arbor, Michigan 48109-1063, USA; M. Urbat, Department of Geophysics, Paleomagnetic Laboratory, University of Utrecht, 3584 CD Utrecht, The Netherlands; G. G. Zuffa, Department of Earth Sciences, University of Bologna, Via Zamboni 67, 40127 Bologna, Italy.

References

100 -

(~

197

Hole 10381

(c)

Fig. 13. Concentration of (a) Li; (b) Na; and (e) K in sediments from Hole 1038I. Conjugate pore fluid concentrations are given in Fig. 10. The dashed lines indicate a basalt intrusion in the sedimentary sequence.

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R . H . JAMES ET AL.

BISCHOFF, J. L. & ROSENBAUER, R. J. 1988. Liquidvapor relations in the critical region of the system NaC1-H20 from 380 to 415°C • A refined determination of the critical point and two-phase boundary of seawater. Geochimica et Cosmochimica Acta, 52, 2121-2126. BUTTEREIELD, D. A., McDuFE, R. E., FRANKLIN,J. & WHEAT, C. R. 1994. Geochemistry of hydrothermal vent fluids from Middle Valley, Juan de Fuca Ridge. In: Mowat, M. J., DAvis, E. E, FISHER, A. T. & SLACK, J. F. (eds), Proceedings of the ODP, Scientific Results. 139, College Station, TX (Ocean Drilling Program), 395-410. CAMPBELL, A. C., GERMAN, C. R., PALMER, M. R., GAMO, T. & EDMOND, J. M. 1994. Chemistry of hydrothermal fluids from the Escanaba Trough, Gorda Ridge. In: MORTON, J. L., ZIERENBERG,R. A. & REISS, C. A. (eds) Geologic, Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern California. U.S. Geological Survey Bulletin, 2022, 201-221. CHAN, L. H., GIESKES,J. M., You, C. F. & EDMOND,J. M. 1994. Lithium isotope geochemistry of sediments and hydrotherrnal fluids of the Guaymas Basin, Gulf of California. Geochimica et Cosmochimica Acta, 58, 4443-4454. CHARLOU,J. L., FOUQUET,Y., DONVAt, J. P., AUZENDE, J. M., JEAN-BAPTISTE, P., STIEVENARD, M. MICHEL, S. 1996. Mineral and gas chemistry of hydrothermal fluids on an ultrafast spreading ridge: East Pacific Rise, 17° to 19°S (Naudur Cruise, 1993): phase separation processes controlled by volcanic and tectonic activity. Journal of Geophysical Research, 101, 15 899-15 919. CLAGUE, D. A. & HOLMES, M. L. 1987. Geology, petrology, and mineral potential of the Gorda Ridge: U.S. Geol. In: SCHOLL,D. W., GRANTZ, A. t~ VEDDER, J. G. (eds) Geology and Resources Potential of the Continental Margin of Western North American Adjacent Ocean Basins-Beaufort Sea to Baja California. Earth Science Series, Circum-Pacific Council for Energy and Mineral Resources, 6, 563-580. DAvis E. A. & FISHER, A. T. 1994. On the nature and consequences of hydrothermal circulation in the Middle Valley sedimented rift: Inferences from geophysical and geochemical observations, Leg 139. In: MOTTL M. J., DAVIS, E. E, FISHER, A. T. & SLACK, J. F. (eds) Proceedings of the ODP, Scientific Results, 139, College Station, TX (Ocean Drilling Program), 695-717. DAVlS, E. E. & VIttINGER, H. 1992. Tectonic and thermal structure of the Middle Valley sedimented rift, Northern Juan de Fuca Ridge. In: Moa--rt, M. J., DAvis, E. E, FISHER,A. T. & SLACK,J. F. (eds) Proceedings of the ODP, Initial Report, 139, College Station, TX (Ocean Drilling Program), 9-41. , MOTTL, M. J., FISHER, A. T. et al. 1992. Proceedings of the ODP, Initial Reports, 139, College Station, TX (Ocean Drilling Program). DUCKWORTH, R. C., FALL1CK,A. E. & RICKARD,D. 1994. Mineralogy and sulfur isotopic composition of the Middle Valley massive sulphide deposit,

northern Juan de Fuca Ridge. In: MOTTL, M. J., DAVIS, E. E, FISHER, A. T. & SLACK,J. F. (eds), Proceedings of the ODP, Scientific Results, 139, College Station, TX (Ocean Drilling Program), 373-385. FOUQUET, Y., WAFIK, A., CAMBON, P., MEVEL, C., MEYER, G. & GENTE, P. 1993. Tectonic setting and geochemical zonation in the Snakepit sulfide deposit (Mid-Atlantic Ridge at 23°N). Economic" Geology, 88, 2018-2036. , ZIERENBERG,R. A., MILLER, D. J. et al. (1998). Proceedings of the ODP, Initial Reports, 169, College Station, TX (Ocean Drilling Program). Fox, C. G. 1990. Consequences of phase separation on the distribution of hydrothermal fluids at ASHES vent field, Axial Volcano, Juan de Fuca Ridge. Journal of Geophysical Research, 95, 12923-12926. GIESKES, J. M., ELDERFIELD, H., LAWRENCE, J. R., JOHNSON, J., MEYERS, B. & CAMPBELL, A. 1982. Geochemistry of interstitial waters and sediments, Leg 64, Gulf of California. In: CURRAY, J. R., MOORE, D. G. et al. (eds) Initial Reports of the Deep Sea Drilling Program. Washington (U.S. Government. Printing Office), 64, 675-694. Ht0MPHRIS, S. E., HERZIG, P. M., MILLER, D. J. et al. 1995. The internal structure of an active sea-floor massive sulfide deposit. Nature, 377, 713-716. JAMES, R. H. & ELDERFIELD, H. 1996. Chemistry of ore-forming fluids and mineral formation rates in an active hydrothermal sulfide deposit on the Mid-Atlantic Ridge. Geology, 24, 1147-1150. JOHNSON, J. W., OELKERS, E. H. & HELGESON, H. C. 1992. SUPCRT92: A software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species, and reactions from 1 to 5000 bar and 0 to 1000°C. Computer Geoscience, 18, 899-947. KARLIN, R., LYre, M. & ZAHN, R. 1992. Carbonate variations in the northeast Pacific during the Late Quaternary. Paleoceanographv, 7, 43-61. & ZtERENBERG,R. A. 1994. Sedimentation and neotectonism in the SESCA area, Escanaba trough, southern Gorda Ridge. In: MORTON, J. L., ZIERENBERG, R. A. & REISS, C. A. (eds), Geologic, Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern California. U.S. Geological Survey Bulletin, 2022, 131-142. KARSTEN, J., HAMMOND,S. R., DAVIS, E. E. & CURRIE, R. G. 1986. Detailed geomorphology of the Endeavour segment of the Juan de Fuca Ridge. Geological Society of America Bulletin, 97, 213 221. KASTNER, M. & GIESKES,J. M. 1976. Interstitial water profiles and sites of diagenetic reactions, Leg 35, DSDP, Bellingshausen Abyssal Plain. Earth and Planetary Scientific Letters, 33, 11-20. , MARTIN, J. B. & ELDERFtELD,H. 1991. Fluids in convergent margins: What do we know about their composition, origin, role in diagenesis and importance for oceanic chemical fluxes? Philosophical Transactions of the Royal Society, London, A, 335, 243-259. KOSK1, R. A., SHANKS,W. C., BOHRSON, W. A. & OSCARSON, R. L. 1988. The composition of -

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DRILLING OF SULPHIDE ORE DEPOSITS: ODP LEG 169 massive sulfide deposits from the sediment-covered floor of Escanaba Trough, Gorda Ridge: Implications for depositional processes. Canadian Mineralogist, 26, 655-674. MAGENHEIM,A. J. & GIESKES,J. M. 1994. Evidence for hydrothermal fluid flow through surficial sediments, Escanaba Trough. In: MORTON, J. L., ZIERENBERG, R. A. &REISS, C. A. (eds) Geologic, Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern California. U.S. Geological Survey Bulletin, 2022, 241255. MORTON, J. L., HOLMES, M. L. & KOSKI, R. A. 1987. Volcanism and massive sulfide formation at a sedimented spreading center, Escanaba Trough, Gorda Ridge, Northeast Pacific Ocean. Geophysical Research Letters, 14, 769-772. - - , ZIERENBERG, R. A. &REISS, C. A. 1994. Geologic, Hydrothermal and Biologic Studies at Escanaba Trough: An Introduction. In: MORTON, J. L., ZIERENBERG, R. A. &REISS, C. A. (eds) Geologic, Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern California. U.S. Geological Survey Bulletin, 2022, 1-18. MOTTL, M. J., DAVIS, E. E., FISHER, A. T. & SLACKJ. F. 1994. Proceedings of the ODP, Scientific Results 139, College Station, TX (Ocean Drilling Program). NORMARK, W. R., GUTIVIACHER,C. E., ZIERENBERG, R. A., WONG, F. L. & ROSENBAUER, R. J. 1994. Sediment fill of Escanaba Trough. In: MORTON, J. L., ZIERENBER6, R. A. &REISS, C. A. (eds) Geologic, Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern California. U.S. Geological Survey Bulletin, 2022, 91-130. ROHR, K. M. M. & SCHMIDT, U. 1994. Seismic structure of the Middle Valley near sites 855858, Leg 139, Juan de Fuca Ridge. In: MOTTL, M. J., DAVIS,E. E, FISHER,A. T. & SLACK,J. F. (eds), Proceedings of the ODP, Scientific Results, 139, College Station, TX (Ocean Drilling Program), 3-17. SEYFRIED, W. E., DIN6, K. & BERNDT, M. E. 1991. Phase equilibria constraints on the chemistry of hot spring fluids at mid-ocean ridges. Geochimica et Cosmochimica Acta, 55, 3559-3580. SHANKS, W. C. & BISCHOEV,J. L. 1977. Ore transport and deposition in the Red Sea geothermal system: a geochemical model. Geochimica et Cosmochimica Acta, 41, 1507-1519.

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Precipitation of hydrothermal sediments on the active TAG mound: implications for ochre formation H. C. G O U L D I N G

t, R. A. M I L L S & R. W. N E S B I T T

School of Ocean and Earth Science, Southampton Oceanography Centre, European Way, Southampton S014 3ZH, UK 1present address." Department of Trade and Industry, 151 Buckingham Palace Road, London, SW1 W 9SS, UK Abstract: Submersible and drilling studies of the active TAG hydrothermal mound (26°N, Mid-Atlantic Ridge) have led to new models of fluid flow and evolution within an active mineral deposit which has implications for Fe-oxide and ochre precipitation. Metalliferous sediments from the top of the hydrothermal mound accumulate from a combination of processes including slumping and oxidation of chimney material and in situ precipitation of low-temperature phases from fluids that percolate through the mound. Geochemical proxies of hydrothermal processes allow identification of the mode of formation of one sediment core from the southeastern periphery of the TAG mound. The Fe rich sediment is capped with a ~5 cm thick kaolinite, illite, chlorite, smectite layer which formed from alteration and replacement of basalt and diagenetic reactions within the hydrothermal sediment. Underlying this layer is a ~10cm thick zone of Mn, Cu, Zn and Pb enrichment which is controlled by the sharp redox gradients in the core. The base of the core is characterized by Mn-poor, Fe-rich oxide that is dominated by goethite, haematite and amorphous Fe oxides equivalent to ochreous and gossan material. Rare earth element (REE) patterns from the different layers within the core allow interpretation of the modes of formation of sediment in the light of existing fluid flow models for TAG. The basal layer is dominated by in situ precipitation of Fe oxide phases from evolved fluids that result from significant withinmound anhydrite precipitation. The REE data for the upper part of the core demonstrate mixing with sea water which provides the oxidizing conditions for Mn precipitation along with Cu, Zn and Pb enrichment from the evolved fluid. Sea water ingress results in higher V/ Fe and P/Fe ratios in the upper part of the core but no enhanced U/Fe ratio. The uppermost clay-rich layer hosts the majority of the REE inventory for the core and the significant positive Eu anomaly indicates recrystallisation of the phyllosilicate phases from the ochreous material during diagenesis. REE data from land-based ochre and gossan sediments demonstrate that the TAG model may be applicable to a wide variety of sites throughout the geological record.

The association of metalliferous sediments with active mid-ocean ridges has been observed since the 1960s with the development of models of seafloor spreading (Bostrom & Peterson 1966, 1969). With the identification of ophiolites as fragments of ancient seafloor (Gass 1968), many of the extensive metalliferous deposits in the geological record could also be attributed to hydrothermal processes occurring at ancient oceanic spreading centres (Robertson 1976; Robertson & Fleet 1986; Robertson & Degnan this volume). Metalliferous sediments have been studied from ancient seafloor settings including the Pindos ophiolite, Greece (Robertson & Varnavas 1993), the Troodos ophiolite, Cyprus (Constantinou & Govett 1972; Robertson 1976; Boyle 1990) and the Semail ophiolite, Oman (Fleet & Robertson 1980; Robertson & Fleet

1986; Karpoff et al. 1988). Hydrothermal sediments in Cyprus include ochre (Fe-rich, Mnpoor sediment), and umber (Fe- and Mn-rich sediment), the former being attributed to sulphide weathering and oxidation (Robertson 1976) and the latter to distal plume fall-out (Boyle 1990). In Oman, the sediments have previously been allocated to several types, namely: sediments associated with gossans, proximal to distal sediments derived from plume fall-out from high-temperature vents, and sediments formed at low-temperature vents or through the interaction of hydrothermal fluid and pelagic sediment (Robertson & Fleet 1986). It should be noted that the metalliferous sediments of Oman do not fall into the distinct ochre and umber categories of Cyprus, but show a complete range of M n and Fe contents.

GOULDING,H. C., MILLS,R. A. & NESBITT,R. W. 1998. Precipitation of hydrothermal sediments on the active TAG mound: implications for ochre formation. In: MILLS,R. A. 8~ HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 201-216

201

202

H.C. GOULDING E T AL.

Black Smoker Complex

# OCHRE ~ . ' SAMPLE ~ro~0~

"3655'

Kremlin Area

3~

,'0

3'0 ;0 s'0

Distance (m)

Fig. 1. Map of the active TAG hydrothermal mound at 26°08'N on the Mid-Atlantic Ridge (adapted from Humphris et al. 1995). Core 2594 position is shown along with the Black Smoker Complex and the Kremlin Area, Core 2598 (Mills et al. 1996) and ochre sample positions (Tivey et al. 1995; Mills & Elderfield 1995a,b). TAG-1 through TAG-5 represent areas of drilling by the Ocean Drilling Program (Humphris et al. 1995). Significant anhydrite was recovered from drilling areas TAG-l, TAG-2 and TAG-5 (Humphris et al. 1995; Mills & Tivey in press).

Metalliferous sediments in modern hydrothermal systems form in three ways. The first process is plume fallout (Heath & Dymond 1977; Dymond 1981; Barrett & Friedrichsen 1982; Barrett et al. 1987; Metz et al. 1988; German et al. 1993; Mills et al. 1993), whereby oxides in the hydrothermally derived plume precipitate and settle to the seafloor under gravity (Dymond & Roth 1988). Other sediments form as mound sulphides, are oxidized by seafloor weathering, and redistributed by mass wasting and slumping (Metz et al. 1988; German et al. 1993; Mills et al. 1993) and subsequent alteration (Dill et al. 1994). The third process is the direct precipitation of oxides from low temperature fluids (Toth 1980; McMurtry et al. 1983; Alt 1988a; Puteanus et al. 1991; Mills et al. 1996). In some cases, these three processes can be distinguished texturally and geochemically. However the recognition of discrete mechanisms of sediment formation is often difficult.

In modern seafloor settings, the term ochre has generally been used to describe indurated red siliceous Fe-oxide material from seamounts and the periphery of sulphide mounds (Alt 1988a; Thompson et al. 1988; Dill et al. 1994; Tivey et al. 1995). Microbial enhancement of Fe-oxide precipitation (Alt 1988a) and diagenetic alteration of secondary mineralization of sulphide debris (Alt 1988b; Dill et al. 1994) have been inferred to contribute to ochre formation. One area where all of these processes are occurring is the Trans-Atlantic Geotraverse (TAG) active hydrothermal mound at 26°N on the MidAtlantic Ridge (MAR).

The TAG Mound, 26°N Mid-Atlantic Ridge The active TAG hydrothermal mound is located at a depth of 3670m on ~100000a crust near the foot of the eastern wall of the median valley within the larger TAG hydrothermal field (Rona

HYDROTHERMAL SEDIMENTS FROM TAG

203

al. 1993). The active mound is circular, metres thick fill depressions in the valley floor approximately 150-200 m in diameter, and rises downslope from inactive deposits to the north of ~ 5 0 m above the surrounding seafloor (Fig. 1). the active mound (Metz et al. 1988). Sediment The approximately 5 Mt sulphide mound surface formation is controlled by both plume fallout is covered entirely by hydrothermal precipitates and mass wasting of mound material (Metz et al. and sediments through which extensive fluid 1988), and the relative inputs from these two flow occurs at a wide variety of temperatures sources can be estimated geochemically (Ger(Thompson et al. 1988; Tivey et al. 1995; man et al. 1993; Mills et al. 1993). In addition, Humphris et al. 1995). Fluids exiting the surface some hydrothermal sediment forms by direct of the mound have been sampled and analysed precipitation at the seafloor in zones of diffuse (Campbell et al. 1988; Edmond et al. 1995; flow (Mills et al. 1996). Gamo et al. 1996; Edmonds et al. 1996; James & The objective of this study is to use the Elderfield 1996; Mills et al. 1996). Comparison mineralogy and geochemistry of seafloor metalof the composition of end-member black smoker liferous sediments to elucidate the mechanisms fluids with a range of lower temperature fluids for formation of Fe-oxide sediments in areas of has allowed estimation of the processes occur- low-temperature hydrothermal flow. The TAG ring within the mound which include large hydrothermal area is a suitable site for such a amounts of anhydrite, pyrite and silica precipi- study as the mechanisms of sediment formation tation coupled with sphalerite dissolution and via plume fallout and mass wasting of sulphide metal remobilisation (Edmond et al. 1995; Tivey debris have been well characterized by previous et al. 1995; James & Elderfield 1996; Mills et al. studies (Scott et al. 1978; Shearme et al. 1983; 1996). It is apparent that sea water entrainment Metz et al. 1988; German et al. 1993; Mills et al. plays a significant role in the formation and 1993). Thus the geochemical signature of lowultimate composition of massive sulphide depos- temperature precipitation of Fe-oxides can be its (Humphris et al. 1995; Mills & Tivey in identified and interpreted within the framework press). of current models of fluid flow and mound The range of sulphide deposits and their formation for TAG. distribution and chemistry have also been used to deduce patterns of fluid flow and mixing within the mound (Tivey et al. 1995). The Geochemical proxies in metalliferous proposed model involves entrainment of sea sediments water into the mound resulting in mixing of sea water and hydrothermal fluid, precipitation of Various geochemical proxies for the different minerals, generation of a more acidic fluid, and processes of hydrothermal sediment formation subsequent metal remobilization (Tivey et al. have been developed through studies of TAG 1995). Subsurface drilling by the Ocean Drilling sediments (see Mills & Elderfield 1995a for Program has extended the sampling to three review). Oxyanions are generally conservative dimensions and a greater understanding of the in sea water and are co-precipitated with Feeffects of sea water entrainment and fluid oxides as they precipitate in the buoyant evolution on the geochemistry and composition hydrothermal plume. Thus As/Fe, P/Fe, V/Fe, of the sulphide deposit (Fig. 1, Humphris et al. U/Fe, Mo/Fe ratios tend to be constant in 1995). hydrothermal plume particulates and in unalThe TAG mound surface is covered in areas tered sediments underlying areas of plume of soft sediment through which diffuse fluids dispersion (Trefry & Metz 1989; Feely et al. percolate (Mills et al. 1996). The mound is 1991; German et al. 1991, 1993; Mills et al. bordered by steep slopes of sulphide debris and 1993). There is some variation in the geochemmetalliferous sediments. Weathering is intense istry of these oxyanions in sea water and As, V and produces gossans of Fe-oxides, secondary and P show much greater particle affinities than sulphides, atacamite and jarosite (Hannington et U and Mo (Turner & Whitfield 1979). Chalcoal. 1988; Herzig et al. 1991; Hannington 1993). phile elements such as Cu, Zn and Pb all form Fe-oxide debris at the base of the TAG mound is sulphide phases that build up as substrate on the coarse grained but becomes finer away from the seafloor without significant interaction with sea mound (Tivey et al. 1995). The metalliferous water. These sulphide phases are then subject to sediments of the surrounding TAG area are seafloor weathering that leads to U enrichment admixed with variable amounts of biogenic on pyrite surfaces (Mills et al. 1994) and mass carbonate ooze and occur as patches covering wasting and transport to sediments (Mills et al. the surrounding pillow basalts (Scott et al. 1978; 1993). The geochemical signature of sediments Shearme et al. 1983). Ponds of sediment several thus derived is distinct due to their elevated et

204

H.C. GOULDING E T AL.

U/Fe ratios and their high chalcophile element content. The rare earth elements (REEs) are potentially the most useful geochemical tool for identifying processes of sediment formation (Ruhlin & Owen 1986; Olivarez & Owen 1989; Mills & Elderfield 1995b). The chemical coherence of the group, coupled with the predictable range in solubility and complexation chemistry with increasing atomic number and decreasing ionic radius, make the REEs ideal tracers in hydrothermal systems which have sharp gradients in physical conditions. The lanthanide group exists predominantly in a trivalent state in natural systems, though Ce 4 + and Eu 2 + can exist under oxidizing and reducing conditions respectively and exhibit anomalous geochemical behaviour. REE data are typically normalized to one component of the system under study to enable interpretation of relative abundances in terms of the processes occurring. Here, data are normalized to chondrite values (Evensen et al. 1978). In sea water, the concentrations of REEs are about a million times less than those of chondrite for the light REE (LREE: La to Eu) and somewhat enriched in the more soluble heavy REEs (HREE: Gd to Lu)(Elderfield 1988). Chondrite-normalized sea water REE patterns display a negative Ce anomaly due to preferential scavenging of relatively insoluble Ce 4+ by suspended particulates (Turner & Whitfield 1979). By contrast, end-member hydrothermal fluids are 1000-10000 times more enriched in REEs and the chondrite-normalized REE patterns show a positive Eu anomaly, LREE enrichment, HREE depletion but no Ce anomaly (e.g. Klinkhammer et al. 1994; Mitra et al. 1994). Hence, the size of the Ce and Eu anomalies and the degree of LREE enrichment can give insights into the relative proportions of sea water and hydrothermal fluid present and any evolution in REE pattern. Mixing of hydrothermal fluids with sea water generates a suite of sulphide and sulphate mineral phases as chimney structures and mound features (Tivey et al. 1995). The REE patterns of such phases are dominated by REEs derived from the hydrothermal fluids, i.e. they exhibit positive Eu anomalies, LREE enrichment and no Ce anomaly (Mills & Elderfield 1995b). Pristine sulphides contain low concentrations of REE due to the incompatibility of the large REE cations with sulphide lattice spacings (Bence 1983). Anhydrite has relatively high levels of REE, because REE cations readily substitute for Ca in the lattice (Morgan & Wandless 1980). Crystallographic control with

continuing anhydrite precipitation leads to evolution of the fluid composition as mineralisation processes fractionate the large LREEs, especially Nd, from the smaller HREEs and discriminate against Eu 2+ uptake (Mills & Elderfield 1995a,b; Humphris 1998). The residual fluid therefore has a higher Lan/Ndn ratios, a lower Ndn/Ybn ratio and a larger Eu anomaly than the original fluid. Anhydrite samples and fluids collected from the Kremlin area both by submersible and drilling, demonstrate the extent of this fluid evolution (Mitra et al. 1994; Mills & Elderfield 1995b; Humphris 1998). Plume precipitation processes are dominated by sea water scavenging of REEs. Particles filtered from the buoyant and neutrally buoyant plume demonstrate the rapid nature of this uptake and dominance of the REE pattern by the sea water signature (German et al. 1990). Similarly, the underlying plume-derived sediments exhibit sea water REE patterns (German et al. 1993; Mills et al. 1993) as do most sediments collected from mid-ocean ridge settings on a world-wide basis (e.g. Ruhlin & Owen 1986; Barrett et al. 1987). Low-temperature fluids collected from TAG exhibit a wide range in REE geochemistry that reflects dilution with sea water and evolution of the fluid composition with subsurface mineralization and dissolution processes (Mitra et al. 1994; Mills & Elderfield 1995b; James & Elderfield 1996). Understanding the fluid evolution requires a detailed knowledge of the fluid reaction pathways through the mound that is now becoming available from combined submersible and ocean drilling studies of the active seafloor deposit (Tivey et al. 1995; Humphris et al. 1995; Mills & Tivey 1998). The aim of this study was to use the suite of geochemical proxies to identify mechanisms of sediment formation within an active low-temperature area of the TAG mound and to compare these data with equivalent deposits from two sites from ophiolite sequences. Ochreous deposits that cap sulphide ore bodies in Cyprus and sediments associated with gossans from Oman are considered to be the closest analogues to the TAG sediment for this study. Sediments from the Skouriotissa ore body, Troodos and the base of the Zuha gossan, Semail ophiolite are compared here to the active TAG deposit.

Sampling and methods A 25 cm long push core from the southeast flank of the active TAG mound was collected by DSRV Alvin from RV A t l a n t i s H in 1993 (Dive

205

HYDROTHERMAL SEDIMENTS FROM TAG Table 1. Major and minor element geochemistry for core 2594

Depth (era)

Si (%)

Ti

A1

(%)

(%)

Mn (%)

Fe (%)

0-1 1-2.5 2.5-4 4-6 6-9 9-12.5 12.5-15 15-17 17-20.5 20.5-22 22-24 24-25

7.35 7.66 7.11 13.8 14.3 7.76 12.5 12.8 6.30 2.99 2.41 1.85

0.096 0.132 0.066 0.072 0.042 0.042 0.048 0.03 0.018 0.018 0.024 0.024

1.13 1.6 0.868 0.821 0.482 0.323 0.323 0.154 0.037 0.169 0.18 0.175

4.8 45.6 7.86 4 1 . 4 15.2 37.1 11.5 3 0 . 8 7.59 35.4 4.79 48.3 0.411 45 0.186 45.6 0.194 56.8 0.139 61.7 0.124 62.9 0.116 63.7

Mg (%)

Ca (%)

P

1.2 1.44 1.53 1.56 1.31 0.96 1.05 0.84 0.456 0.498 0.48 0.348

1.21 1.11 1.07 0.764 0.536 0.586 0.479 0.336 0.271 0.343 0.336 0.264

0.729 0.703 0.563 0.266 0.201 0.306 0.227 0.17 0.201 0.293 0.214 0.192

(%)

V U Pb Cu Zn (ppm) (ppm) (ppm) (ppm) (ppm)

93 87 79 50 41 29 11 5.0 5.0 7.0 4.0 4.0

5.0 5.0 5.0 5.0 7.0 10 10 10 13 15 1! 11

1810 1290 ll90 898 675 457 313 282 274 198 196 349

2300 2080 2450 2250 2440 2370 1160 610 518 332 240 239

1690 1170 1360 666 428 374 240 159 165 176 180 504

Precision ofSi, A1, Fe, Mn, Ca and P data is < 5% (2a), and 5-10% (2a) for Ti and Mg (all determined by XRF). Precision of V, U, Pb, Cu and Zn data is better than 3% (2or) (determined by XRF, determined by ICP-MS).

2594), and stored at 4°C. This core was collected from 3660m depth , - d 0 m SE of the Kremlin white smoker chimney area and ~70 m E of core 2598 described by Mills et al. (1996) (Fig. 1). The core was sectioned longitudinally, and subsamples were taken at 1 - 4 c m intervals to represent the major units defined by colour and texture differences. Samples were dried in an oven at 60°C, then ground to a fine powder with an agate pestle and mortar. For X R D analyses, two slides were prepared from each specimen to allow for b o t h bulk m i n e r a l o g y and clay determinations. For clay mineral determination, the Fe in the sample was removed by the sodium citrate bicarbonate dithionate method (Jackson 1969), before the < 2 #m fraction was separated by standard procedures. Major and trace elements were determined by X-ray fluorescence (XRF). Glass beads were prepared by mixing the ignited sample with dilithium tetraborate fluxing agent in ratios of 1: 10, or 1:20 for samples with high Fe or Mn content, to allow a homogeneous mix to be obtained. Trace elements, including the REEs, were measured by Inductively Coupled Plasma Mass Spectrometry (ICP-MS) (Barrat & Nesbitt 1996; Barrat et al. 1996). All analytical procedures were carried out at the Southampton Oceanography Centre (SOC). Three sediment samples were collected from the base of the Zuha gossan, Semail ophiolite, Oman (Karpoff et al. 1988; Goulding 1998). These samples (A, B1 and B2) were analysed with the Core 2594 samples to allow direct comparison of sediments from the two settings.

Results

Mineralogy

The TAG sediments consist of unconsolidated amorphous Fe-oxides and oxyhydroxides (haematite and goethite) with small amounts of quartz and carbonate present throughout the core. Minor smectite occurs in all samples, illite occurs throughout most of the core and kaolinite and chlorite are present within the top ~5 cm of the core. Minor pyrite occurs throughout the core and chalcopyrite is present in the basal core sample. Minor atacamite is present throughout the core, with higher abundances in the upper part of the core. Centimetre-scale horizontal colour banding is due to changes in mineralogy, with dark brown to reddish colours arising from dominance of haematite, and orange colours due to the presence of goethite. Two centimetre-size fragments of basalt were present in the upper portion of the core. The Z u h a gossan sediment samples are dominated by quartz, haematite, maghaematite, epidote, chlorite, illite, kaolinite and smectite with minor pyrite and calcite. These sediments correspond to the deposits described by Karpoff et al. (1988). Geochemistry

The major and trace element composition of the T A G sediment core is given in Table 1. Downcore variations in some key elements that are indicative of the processes controlling metal

206

H.C. GOULDING E T A L .

(a)

(b)

Mn (%) 0

0

5

P/Fe

to

~s

20

0

0.oo5

0.01

0.015

0.02

1<

2O

25 0

25 10

20

30

40

50

60

t

70

f .....

lxl0 ~

Fe and Si (%)

2xlff'

J

3x10"

V/Fe and U/Fe

(e)

(d) AI/Ti and Nd (ppm) 0

5

O

10

,

, ..........

5

.fi

15

Cu (ppm) 20

[

,

5

,

A

~

0

0

1000

2000

3000

40OO

Zt

tW

15 . Al/ri "'-..

N

25 0

~-~

I I

20

I

0.5 1 1.5 A1 and Ti (%),

2

25 0

|

,,

Zn and Pb (ppm)

Fig. 2. Down-core variation in (a) Mn, Fe and Si, (b) P/Fe, V/Fe and U/Fe ratios, (e) Nd, A1, Ti and A1/Ti ratio, (d) Cu, Pb and Zn for Core 2594. Note Ti concentration data have been scaled by a factor of 10 to allow visual comparison with A1 data. TAG plume particulate values of P/Fe, V/Fe and U/Fe are 6x10 -z, 4.1x 10-5 and 2.0x 10-5 respectively (Trefry & Metz 1989; Feely et al. 1991; Mills et al. 1994). The plume V/Fe ratio is shown as the dashed line in (b). Mid-ocean ridge basalt (MORB) A1/Ti value of 13.5 is depicted as the dashed line in (c).

distribution are shown in Figs l(a to d). The main c o m p o n e n t of the T A G sediments is Fe, p r e s e n t as oxides a n d m i n o r pyrite, with significant M n and Si present. The Fe and Si levels are in the range previously observed for T A G sediments from the active m o u n d area ( G e r m a n e t al. 1993; Mills e t al. 1996). M n levels exceed 15% in the upper part of the core; this extent of enrichment has not been previously

observed in proximal T A G sediments. There is a general negative correlation between Fe and Si content, whereas there is no significant correlation of Fe with M n content (Fig. 2a). Phosphorus and U levels are low compared with previous studies of h y d r o t h e r m a l sediments. There is a positive correlation between U and Fe content down-core (r 2= 0.76). The U / F e ratio of sediments from this core is ~ 2 . 1 x 1 0 -5

14.3 14.9 10.7 8.83 5.27 4.39 3.17 0.643 0.399 0.389 0.371 0.365

0-1 1-2.5 2.544-6 6-9 9 12.5 12.5-15 15-17 17-20.5 20.5-22 22-24 24-25

24.4 26.6 18.3 16.5 11.8 9.04 6.96 1.94 1.29 1.27 0.955 0.943

Ce (ppm)

3.38 3.98 2.65 2.49 1.48 1.28 0.997 0.306 0.262 0.233 0.154 0.157

Pr (ppm)

15.4 18.4 12.4 11.5 6.99 5.95 4.91 1.84 1.45 1.25 0.903 0.912

Nd (ppm)

3.33 3.9 2.6 2.48 1.53 1.26 1.18 0.594 0.459 0.413 0.258 0.266

Sm (ppm) 2.98 3.41 2.41 2.12 1.83 1.67 1.4 1.39 1.48 1.36 1.01 0.897

Eu (ppm) 3.3 3.68 2.54 2.25 1.33 1.21 0.954 0.417 0.313 0.27 0.208 0.204

Gd (ppm) 0.499 0.538 0.353 0.326 0.188 0.167 0.138 0.05 0.04 0.039 0.03 0.03

Tb (ppm) 2.94 3.1 2.02 1.83 1.04 0.906 0.743 0.238 0.178 0.181 0.141 0.14

Dy (ppm) 0.654 0.663 0.447 0.384 0.217 0.19 0.141 0.042 0.029 0.028 0.023 0.022

Ho (ppm) 1.87 1.87 1.24 1.08 0.6 0.503 0.373 0.108 0.069 0.07 0.058 0.057

Er (ppm) 0.255 0.249 0.168 0.148 0.083 0.067 0.048 0.014 0.009 0.009 0.008 0.008

Tm (ppm) 1.570 1.560 1.02 0.922 0.514 0.411 0.321 0.090 0.050 0.050 0.048 0.046

Yb (ppm) 0.262 0.252 0.169 0.150 0.083 0.065 0.048 0.013 0.008 0.008 0.007 0.008

Lu (ppm) 0.805 0.802 0.792 0.820 0.984 0.891 0.918 1.02 0.900 0.963 0.938 0.923

Ce/ Ce*

2.75 2.74 2.87 2.73 3.88 4.13 3.96 8.21 11.4 11.8 13.1 11.5

Eu/ Eu*

Precision for R E E measurements is better than 3% (2o-), Ce anomaly is expressed as Cen/Ce* where Ce* = (Lan + Pr,/2), Eu anomaly is expressed as Eun/Eu*, where Eu* = (Smn + Gdn/2) and L R E E enrichment is expressed as Ndn/Ybn ratio. All ratios are calculated using chondrite normalized values (Evensen et al. 1978).

La (ppm)

Depth (cm)

Table 2. R E E data and anomalies f o r core 2594

3.42 4.11 4.24 4.35 4.74 5.04 5.33 7.12 10.1 8.71 6.55 6.91

Ndn/ Ybn

208

H. C. GOULDING E T AL. 100 '

_

North Atlantic

DeepWater

(* 10 7)

Top

I 1 Base 0.1

La Ce

P r Nd

Sm Eu Gd Tb D y Ho E r T m Yb Lu

Fig. 3. Chondrite normalized REE data for Core 2594. The core top data exhibit the highest chondrite normalized values and the core base the lowest. The shaded area encompassed by the dashed line represents data from TAG ochres (Fig. 1; Mills & Elderfield 1995) and the dashed line represents sea water REE data scaled by 107 (Mitra et al. 1994). Table 3. Geochemical data for three sediment samples from the base of the Zuha gossan, Semail Ophiolite, Oman Element

sample A

sample B 1

sample B2

Si (%) Ti (%) A1 (%) Mn (%) Fe (%) Mg (%) Ca (%) P (%) V (ppm) Pb (ppm) Cu (ppm) Zn (ppm) La (ppm) Ce (ppm) Pr (ppm) Nd (ppm) Sm (ppm) Eu (ppm) Gd (ppm) Tb (ppm) Dy (ppm) Ho (ppm) Er (ppm) Tm (ppm) Yb (ppm) Lu (ppm) Ce/Ce* Eu/Eu* Ndn/Ybn

29.6 0.072 1.18 5.09 8.19 0.21 9.53 0.36 122 59.3 241 151 29.0 11.6 4.79 21.0 4.14 1.88 5.01 0.783 4.72 1.06 2.87 0.373 2.35 0.358 0.214 1.28 3.11

27.0 0.29 0.609 0.194 13.4 0.09 2.74 0.33 238 32.2 2150 269 30.9 20.3 4.88 17.6 2.63 1.37 2.26 0.268 1.48 0.358 1.48 0.269 2.16 0.389 0.356 1.70 2.84

21.4 0.19 0.307 0.38 33.6 0.10 2.09 0.26 367 6.9 1810 286 27.3 11.6 3.87 13.6 1.67 0.727 1.32 0.178 1.17 0.314 1.28 0.231 1.84 0.328 0.238 1.46 2.58

All data obtained by XRF and ICP-MS, precision as for Tables I and 2.

(by weight) (Fig. 2b) which is comparable to previous studies of plume-derived and Fe-oxide metalliferous sediments from T A G (2x10-5; Mills et al. 1994). The P/Fe ratio is low compared with TAG plume particles (0.063; Feely et al. 1991) and decreases significantly down-core. The V/Fe ratio is high at the core top and decreases to levels below those observed for TAG plume particles (4.1x10-5; Trefry & Metz 1989) in the lower part of the core. Aluminium shows a general decrease downcore which is also shown by the Ti and REE data (Fig. 2c). There is a close positive correlation between AI and Ti (r 2= 0.97), and each of these elements with the REEs. Aluminium, Ti and REE levels are in the range previously observed with these types of sediments (Metz et al. 1988; German et al. 1993; Mills et al. 1993). The A1/Ti ratio is close to N - M O R B values in the upper part of the core, decreases to a minimum at ~20 cm and increases again at the core base (Fig. 2c). The chalcophile elements, Cu, Zn and Pb are all enriched within the upper portions of this short core (Zn and Pb contents exceed 1000 ppm, Cu content exceeds 2000 ppm; Fig. 2d). These levels are all high compared with previous studies of TAG sediments (Shearme et al. 1983; Metz et al. 1988; German et al. 1993) and decrease significantly down-core. Ca and Mg levels are low throughout the core ( < 2 % throughout; Table 1). REE data are shown in Table 2 and Fig. 3. The chondrite-normalized REE patterns from the T A G core all display significant positive Eu

HYDROTHERMAL SEDIMENTS FROM TAG anomalies (Eu/Eu*=2.74-11.5), negligible Ce anomalies (Ce/Ce*=0.79-1.0) and LREE enrichment (Ndn/Ybn = 3.42-10.1). The size of the Eu anomaly increases systematically down-core, as does the LREE enrichment. The Ce anomaly is only apparent in the upper core (Fig. 3). The composition of three metalliferous sediment samples from the base of the Zuha gossan in Oman is given in Table 3. The geochemistry of these samples is similar to the range of values shown by the TAG sediment. In general, the gossan sediment is A1 and Ti poor, Fe, Cu and Zn and Mn enriched. Ca levels are significantly higher than for TAG sediment and this is attributed to the higher calcite component within the gossan, calcite replaces biogenic silica through much of the sequence. The REE levels are comparable to the TAG core and exhibit distinct Eu anomalies (Eu/Eu* = 1.2-1.7). These sediments also exhibit significant Ce anomalies (Ce/Ce*=0.21-0.36) and some LREE enrichment (Ndn/Ybn = 2.6-3.1).

Discussion The TAG sediment core is similar in mineralogy to ochreous phases from the active TAG mound (Tivey et al. 1995) and from the Skouriotissa deposit in Cyprus (Boyle 1990; Herzig et al. 1991). The atacamite is a result of the supergene enrichment of Cu within the TAG mound and stabilization of this phase by the overlying sea water (Hannington 1993) and is not observed in any ophiolite sequence. The presence of basalt fragments within the upper core is consistent with observations of highly altered basalt clasts at shallow stratigraphic levels in the TAG mound (Humphris et al. 1995; Honnorez et al. 1998). While the mechanism for basalt emplacement atop the active mound is not clear, a model of near complete replacement of the volcanic mound in the southern periphery is consistent with the observation of a ~5 cm thick surficial clay enrichment observed in Core 2594. The clay assemblage is inferred to be the residue of basaltic replacement and weathering. Mg-rich chlorite has been observed under the SE mound (TAG-2 area Fig. 1) and is infered to form via reaction with heated sea water at temperatures from 250-370°C (Honnorez et al. in press). Chlorite is observed in Core 2594 throughout the upper ~5 cm clay rich layer and is derived from basaltic alteration. Previously, surficial clay enrichments (smectite, illite, kaolinite) have been observed in a core from the southern periphery of the mound approximately 70 m due west of Core 2594 (Fig. 1; Mills et al. 1996). In

209

the western core (2598) the clay alteration products are inferred to form at ~80°C (Mills et al. 1996). A similar alteration assemblage is seen in Core 2594 (smectite, illite, kaolinite) emphasizing the importance of in situ lowtemperature mineralization (Alt & Jiang 1991; McMurtry et al. 1983) in this southern region of the mound. Core 2594 is dominated by Fe and Si phases with significant Mn in the upper core (Fig. 2a). Fe levels are extremely high in the lower core (> 60%) as a consequence of the presence of significant goethite phases with minor pyrite. Dilution of the Fe-oxide phases with Fe-silicate material (smectite) and quartz leads to the apparent anti-correlation of Fe with Si downcore. The Fe-oxide layers tend to be finer grained whereas the silicate material is coarser and more permeable and therefore the mineralogy has implications for fluid percolation and channelling within the mound. Ca levels are low throughout the core and X R D data suggest only minor calcite is present. Comparison with published data from other cores collected from the active mound (German et al. 1993; Mills et al. 1996) suggests that this material is biogenic and derived from sedimentation from the overlying water column. There must be a small pelagic clay component associated with the biogenic sedimentation, though this will be insignificant compared with the hydrothermal processes that predominate in this core (Metz et al. 1988). Hydrothermal sediment collected from the surface of an active mound can potentially be derived from a mixture of plume fall-out, sulphide weathering and low-temperature precipitation of Fe-oxides from fluids percolating through the upper mound as discussed earlier. The lack of correlation between Fe and either P or V rules out unmodified plume fall-out as a significant component of this sediment (Trefry & Metz 1989). Sedimentary P/Fe ratios are low compared with overlying plume particles whereas V/Fe ratios are high in the upper part of the core relative to plume particles and low in the lower part of the core (Fig. 2b). The U:Fe ratio is similar to the U:Fe ratio of a wide range of TAG sediments (Mills et al. 1994) including plume derived sediments and fully oxidised sulphidic debris. The U:Fe ratio of 2.1x10 s suggests that any significant sulphide debris within the core has been fully oxidized. XRD data back up this inference and demonstrates that there is only minor pyrite throughout the core. Any original sulphide debris has been replaced during low-temperature alteration of the sediment pile. This alteration has led to the

210

H.C. GOULDING E T AL.

high V/Fe and P/Fe ratios observed in the upper core as sea water V and P diffuses into the upper sediment and is precipitated during secondary mineralization processes. The decoupling of V and P from U is possible as V and P exhibit far greater particle affinity than U in sea water (Trefry & Metz 1989; Feely et al. 1991), whereas U requires reductive fixation at Eh levels below the ferric/ferrous transition (Thomson et al. 1993). Mn is strongly decoupled from Fe precipitation within seafloor hydrothermal systems, most Mn precipitation is associated with low-temperature oxides at some distance from the active high-temperature venting (Scott et al. 1978; Thompson et al. 1985). In Core 2594 there is a broad enrichment of Mn in the upper core sampled ~ 100 m from the black smoker complex at TAG. Sedimentary Mn enrichments have been attributed to transport of Mn oxide material via slumping processes (Metz et al. 1988; Mills et al. 1993) and surficial enrichment of Mn has been attributed to fractionation of this element into distal plume fall-out (Shearme et al. 1983). Diffuse flow fluids at TAG are enriched in Mn (James & Elderfield 1996; Mills et al. 1996) as they represent high-temperature fluids diluted with entrained sea water. Mn precipitation will occur in the upper core as the upwardly diffusing fluid mixes with oxidizing sea water if the Eh conditions are high enough to instigate Mn oxidation. The relatively impermeable clay mineral cap may produce the steep redox gradients that allow Mn precipitation in the upper part of the core. Precipitation of fresh Mn-oxide phases has important implications for the geochemical budgets of many other metals in solution because of its potential to scavenge other metals from solution (Koeppenkastroop & De Carlo 1992). The observation of diagenetic Mn enrichment atop the southern periphery of the active mound demonstrates the extreme heterogeneity in reaction pathways within a single active sulphide deposit and the steep redox gradients in this area. The close correlation between A1, Ti and Nd content (Fig. 2c) demonstrates that the REEs are largely hosted in the alumino-silicate phases present within this core. Kaolinite and chlorite contents are high in the upper ~5 cm of the core and are responsible for the larger A1, Ti and Nd contents observed. Mg-rich chlorite has been recovered from the SE margins of the mound (Honnorez et al. 1998) and the upper part of the core is similarly enriched in Mg (up to 1.6% Mg; Table 1). The major and trace element data for Core 2594 allow identification of the general redox

conditions of sediment formation. The extremely low Mn content of the lower portion of the core implies that Fe-oxide precipitation is initiated at Eh conditions high enough for ferric iron deposition without Mn precipitation. The kinetics of Mn oxidation are slow but the Eh conditions of the upper core allow extensive Mn precipitation along with Cu, Pb, Zn, V and P enrichment (Fig. 2b & d). The chalcophile elements Cu, Zn and Pb are all enriched in the upper core (Fig 2d). Cu is mainly present as atacamite and such supergene enrichment is common at TAG (Hannington 1993). Cu is transported in reducing acidic fluids, complexed by the C1 content of the fluid. Cu precipitation is initiated on mixing with oxidized sea water within the sediment column and atacamite precipitation is high in the upper core and associated with MnO2 precipitation. Pb and Zn enrichment is inferred to occur via a similar mechanism of transport in fluids at low Eh and precipitation within the zone of sharp redox gradient. The close proximity of this core to the Kremlin area (Fig. 1) suggests that many of the processes inferred to control white smoker fluid composition (e.g. Edmond et al. 1995; Tivey et al. 1995; Mills & Elderfield 1995b) are acting here. These processes are common on larger scales within land-based deposits and act to zone refine the resultant ore body (Franklin et al. 1981). The REE patterns are all dominated by a large Eu anomaly which increases down-core. The Eu anomaly for the samples at the base of the core exceeds that of the black smoker fluid end-member (Eu/Eu *= 9.2; Fig. 4a; Mitra et al. 1994) and any pyrite or oxidised sulphide phase collected at TAG (Eu/Eu *= 2-6; Mills & Elderfield 1995b). These patterns are however very similar to those reported for TAG mound ochres (Fig. 3; Mills & Elderfield 1995b) and to lowtemperature derived Fe-oxide sediments (Mills et al. 1996), and a common mechanism for formation is inferred. The four components that define the REE patterns for Core 2594 are the Ce anomaly (Ce/Ce*), the Eu anomaly (Eu/Eu*) and the extent of LREE fractionation (Lan/Ndn) and L R E E / H R E E fractionation (Ndn/Ybn). None of the REE data for Core 2594 fall along the sea water-black smoker mixing lines for these parameters (Figs 3 a & b); therefore these sediments cannot be generated by direct mixing of sea water and black smoker hydrothermal fluid. Oxidation of sulphide phases can also be ruled out since primary sulphides at TAG exhibit extremely low REE contents (Mills & Elderfield 1995a,b) and a mechanism which increases the size of the Eu anomaly and

HYDROTHERMAL SEDIMENTS FROM TAG (a)

Core

•

•e

Base

TAG fluid

evolu.on.. ";

9

ta

B$

6 • .•"

Core Top ~ •

~JSca

s

I

0.2

0.4

, l

0.8

0.6

water i~qu~ee J

1

1.2

Ce/Ce* (b) TAG mound e

10

Core •

fluid

BS

8

4

"

•

Core Top 2 0

"'"" . . . . . SW L

0

0.5

I 1

I ...... 1.5

I. . . . . . 2

I 2,5

LaJNd

Fig. 4. Co-variation in (a) Eu anomaly and Ce

anomaly and (b) Ndn/Ybn and Lan/Ndn. The dashed line represents incremental mixing of TAG black smoker fluid with sea water, BS = TAG black smoker fluid and SW = North Atlantic Deep Water (data from Mitra et al. 1994). The arrows represent evolution of REE patterns with progressive sea water influence or fluid evolution within the mound. Open circles in (a) are Oman metalliferous sediment data with sea water REE patterns from Goulding (1998).

fractionates the L R E E from the H R E E is required to explain the data. Insights from recent drilling of the TAG mound (Humphris et al. 1995) and geochemical studies of anhydrite phases sampled subsurface (Humphris 1998; Mills & Tivey in press) allow reinterpretation of the source of REEs to these sediments and ochres. The only significant sink for REEs during primary sulphide and sulphate formation at TAG is anhydrite (Mills & Elderfield 1995b). Estimates of the TAG anhydrite inventory are of the order of 2x 104 m 3 (Mills & Tivey in press) and precipitation of large volumes of anhydrite can potentially strongly modify the REE content of the residual fluid.

211

Subsequent dissolution of this anhydrite at temperatures < 150°C imparts the trapped REEs into solution and again can influence the REE budget for the mineral deposit. Precipitates of Fe-oxides forming in situ will scavenge REEs from solution. The middle REEs are preferentially scavenged from solutions where the REEs are CO3-complexed such as sea water (Elderfield 1988). LREEs form more stable C1 complexes than the HREEs in acidic solution such as hydrothermal fluids (Wood 1990) and Fe-oxides precipitating from such solutions would be H R E E enriched. Solution complexation will not affect the Eu and Ce anomaly, which should therefore reflect redox state of the parental fluid. Europium will be present as Eu 3+ at temperatures below 250°C and will exhibit trivalent behaviour (Sverjensky 1984) and Ce will be stable as Ce 4+ in fluids with Eh values inferred here from Mn and Fe geochemistry (Turner & Whitfield 1979). The REE inventory of the fluid mixture is scavenged onto the Fe-oxide substrate with some fractionation of the LREEs from the HREEs depending on pH, whereas the size of the Eu anomaly will reflect fluid evolution (Koeppenkastroop & De Carlo 1992). The ochreous material from the base of the core precipitates from a fluid that has a large Eu anomaly and no Ce anomaly (Fig. 4a) that has evolved from large amounts of anhydrite precipitation. However, the low Lan/Nd, and high Ndn/Ybn ratios of these samples (Fig. 4b) cannot be generated by anyhydrite precipitation. Instead, solution complexation at the low Eh and low pH conditions inferred for the core base must fractionate the REEs to produce the observed patterns. The REE enrichment in the upper core is accompanied by a reduction in the size of the Eu anomaly (Eu/Eu*=0.5) and g e n e r a t i o n of a small Ce a n o m a l y (Ce/ Ce* = 0.27) (sea water influence in Fig. 4a). This is consistent with the downwards diffusive penetration of sea water and mixing in the upper sediment core. The close correlation of A1 with REEs suggests that the latter are associated with the phyllosilicate fraction of the sediment and that the diagenetic overprint of the ochre occurs mainly through clay mineral formation. The TAG sediments and ochres at the southern mound periphery precipitate from fluids that have evolved within the mound to exhibit REE patterns that have large positive Eu anomalies (Fig. 1; Mills & Elderfield 1995a,b; Mills et aL 1996). These fluids transport significant metals (James & Elderfleld 1996; Mills et aL 1996) which precipitate in the sharp redox gradients of the upper part of the sediment core. The high

212

H.C. GOULDING ET AL. 100

-

10

1

0.1

I

I

I

I

I

La Ce P r Nd

I

I

I

i~

!

l-"

!

i_.-l-']

Sm Eu Gd Tb Dy Ho E r Tm Yb Lu

Fig. 5. Comparison of the range of chondrite normalized REE patterns for Core 2594 with data for Oman gossans (this study), Skouriotissa ochres (Herzig et al. 1991) and Southern Explorer Ridge gossan residues (Barrett et al. 1990). These ochre and gossan sediment data represent examples of TAG type REE patterns in the geological record.

fluid fluxes are responsible for extensive basalt alteration and secondary mineralization. There is no evidence for sulphide oxidation being a significant process in ochre formation at TAG. Comparison ochres

o f sea f l o o r s e d i m e n t s with

a n d gossan

sediments from

the

geological record

REE data from metalliferous sediments sampled on the flanks of the East Pacific Rise (Barrett & Friedrechsen 1982; Barrett et al. 1987; Olivarez & Owen 1989), Pacific seamounts (Alt 1988a,b) and ophiolite-hosted metalliferous sediments (Robertson & Fleet 1986; Goulding 1998) show sea water derived REE patterns with variable Ce anomalies, and no Eu anomaly. REE anomalies from Semail ophiolite metalliferous sediments (Goulding 1998) are shown in Fig. 4a to demonstrate the clear distinction between sea water-derived REE patterns and the ochreous material discussed here. Any original hydrothermal signature derived from parental fluids has been subsequently overprinted by scavenging of sea water REEs during sediment accumulation and sulphide oxidation and subsequent emplacement and obduction (Robertson & Fleet 1976). Metal-enriched sediments from the Atlantis II Deep in the Red Sea show REE patterns with LREE enrichment, large Eu anomalies (Eu/Eu*

up to 3.45) and no Ce anomaly (Cocherie et al. 1994). These sediments are composed of a complex mixture of hydrothermal minerals including anhydrite, Fe-oxides and clays and the REE patterns observed are inferred to reflect that of the parental hydrothermal fluid with little fluid evolution during percolation through the sediment pile (Cocherie et al. 1994). Similar REE patterns that are inferred to reflect the REE pattern of the parental high-temperature fluid have been observed in various seafloor sulphide-sulphate deposits (Alt 1988b; Barrett et al. 1990). Fe-oxide deposits with fractionated REE patterns and significant positive Eu anomalies have only been observed from a few locations and data are compiled in Fig. 5. Three sites exhibit similar types of REE behaviour and mineralization to the TAG sediment: Skouriotissa ochres, Cyprus (Herzig et al. 1991), sediments from the base of the Zuha gossan, Oman (this study) and Southern Explorer Ridge gossans (Barrett et al. 1990). The Skouriotissa ochres are associated with a TAG-type ore body of equivalent size (~6 Mtonne) though the tectonic and geological setting for Troodos was supra-subduction rather than mid-ocean ridge (Pearce & Cann 1973; Pearce et al. 1985). The Zuha gossan overlies disseminated sulphide and the gossan sediments occur at the base of the weathered cap (Karpoff et al. 1988; Goulding 1998). The Southern Explorer Ridge gossans are

HYDROTHERMAL SEDIMENTS FROM TAG unconsolidated metalliferous sediment overlying the sediment hosted sulphide-sulphate deposits (Barrett et al. 1990). The interbedded ochreous samples from the Skouriotissa ore body in Cyprus exhibit significant REE fractionation. These data show no significant LREE enrichment but the Eu anomaly is large (Fig. 5). The alteration assemblage observed in Skouriotissa, namely atacamite, jarosite and quartz, is very similar to that observed in TAG ochres (Tivey et al. 1995) and a common mode of formation is inferred (Herzig et al. 1991). The Zuha gossan sediments from the Semail ophiolite of Oman are dominated by biogenic (radiolaria) and hydrothermal inputs with little detrital influence on the mineralogy and geochemistry (Karpoff et al. 1988). These sediment samples show positive Eu anomalies (Eu/Eu* up to 1.7), LREE enrichment (Ndn/Ybn up to 3.0) and HREE fractionation that is attributed to fluid evolution during formation. The shape of the REE pattern is similar to that observed for gossan material from the Southern Explorer Ridge (Fig. 5; Barrett et al. 1990). The data compiled in Fig. 5 demonstrate that while most metalliferous sediments from the geological record do not record the fractionated REE patterns determined during formation, some samples of banded ochreous or sedimentary gossan material in the close proximity to sulphide deposits retain their REE patterns obtained during formation processes akin to that occurring at TAG.

Conclusions While most metalliferous sediments sampled to date exhibit sea water-dominated REE patterns, sediments that are derived from low-temperature precipitation of Fe-oxides in association with sulphide mineralization exhibit fractionated REE patterns. Recent studies of fluid circulation within the active TAG mound (Mills & Tivey in press; Humphris 1998) have led to development of models of fluid evolution with implications for the geochemistry of the deposit. Entrainment of sea water into the permeable sulphide mound leads to precipitation of anhydrite from sea water derived sulphate and Ca sourced from both the vent fluids and the sea water (Mills & Elderfield 1995a,b; Mills & Tivey in press). REEs preferentially partition into the anhydrite phase, the larger LREE readily replacing Ca in the lattice structure (Morgan & Wandless 1980; Mills & Elderfield 1995a,b; Humphris 1998) leaving behind a fluid that has a larger Eu anomaly. When the evolved fluid mixes with sea

213

water, Fe oxides are precipitated directly from solution as ochreous and oxide deposits on the southern periphery of the mound (Tivey et al. 1995; Mills et al. 1996). The REE signature of these deposits provides evidence for the fluid pathways and fluid complexation. The ochreous and metalliferous sediment deposit (Tivey et al. 1995; Mills et al. 1996) at the southern periphery of the TAG mound represents the current major zone of low-temperature deposition at TAG. The fluid flow through these sediments also results in metal mobilization and remineralization leading to Mn, Cu, Zn and Pb enrichment of the upper part of the core. This mineralization sequence provides evidence for sharp Eh gradients in the sediment core between the lower zone of Fe-oxide stability and the upper zone of MnO2 stability. This gradient in physico-chemical conditions ultimately controls the replacement mineralization in the upper mound. The upper layer of clay mineral alteration assemblage demonstrates the importance of basalt alteration and replacement in mound formation and the extensive low-temperature fluid flow alteration occurring. Sea water ingress into the upper core is evident in the development of a negative Ce anomaly in the upper part of the core and the enhanced V/Fe and P/Fe ratios. We have demonstrated that the observations and models arising from work at TAG could help elucidate mechanisms for mineralization and alteration at sites throughout the geological record if sampling was carried out on a scale similar to that on the seafloor. The scarcity of good quality REE data from these sites adjacent to sulphide ore bodies requires rectification in the light of this work at TAG. We would like to thank A. Milton for assistance with ICP-MS analyses, I. W. Croudace for XRF analyses and J. Fisher for XRD analyses. A. Nimmo Smith and K. Davies are thanked for their help with preparation of Figures. P. A Rona is gratefully acknowledged for providing Alvin dive time at TAG. RAM would like to thank the captain, officersand crew of the RV Atlantis H and DSRV Alvin during leg 129-5 for all their assistance with sample collection. HCG is grateful to R. Koski, A. Galley, H. Gibson and P. Wipplinger for assistance in the field in Oman. The manuscript was greatly improved following comments from three anonymous reviewers. This research was funded by NERC Research Studentship GT4/94/298/G and grant BRIDGE 21. RWN acknowledges grants from the EU.

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Significance of modern and ancient oceanic Mn-rich hydrothermal sediments, exemplified by Jurassic Mn-cherts from Southern Greece A. R O B E R T S O N 1 & P. D E G N A N 2

aDepartment o f Geology and Geophysics, University o f Edinburgh, Edinburgh EH9 3JW, UK 2UK Nirex Limited, Currie Avenue, Harwell, Didcot O X l l ORH, UK Abstract: This paper considers the occurrence of Mn-oxide-rich hydrothermal deposits in

the modern oceans and presents a case history of Jurassic Mn-cherts from a Tethyan area in Southern Greece. Manganese-oxide hydrothermal sediments are known from a wide variety of oceanic settings, including both on and off the axis of spreading ridges, in back-arc basins, arc-related areas, rifts and in small ocean basins. These Mn-rich sediments differ markedly from the chemical composition of better known metalliferous deposits associated with hightemperature black smokers (i.e. Fe-rich ochres and Fe-Mn-rich umbers). The ancient example discussed in detail here is that of Late Jurassic Mn-rich cherts spatially associated with basic-intermediate composition volcanics within sutured oceanic units in the Peloponnese, S Greece. The Mn enrichment there is interpreted as related to hydrothermal activity associated with Late Jurassic off-axis volcanism, rather than to a spreading ridge source. The Mn is strongly fractionated from Fe and is concentrated in thin beds and laminations, interpreted as of primary depositional origin. The silica is assumed to be mainly biogenic, as reflected in an association with red radiolarian cherts. The source hydrothermal activity is likely to have been of low-temperature type. Similar Mn-oxide deposits occur elsewhere in the Mesozoic Tethyan area , including N Greece, SW Turkey, Cyprus and Oman, and in other orogenic belts (e.g. in California). There are also metamorphosed counterparts of Mn-rich hydrothermal deposits that include metalliferous pelites and cherts; e.g. Cyclades, S Greece and Franciscan Complex, California. The record of deep-sea hydrothermal deposits exposed on land following tectonic emplacement is biased towards off-axis Mn-deposits within abyssal sediments, since axial spreading ridges (including massive sulphides, umbers and ochres) are rarely preserved, owing to subduction.

Black smokers that precipitate massive sulphides in the oceans and their ancient counterparts associated with ophiolites (e.g. in Cyprus; Oudin & Constantinou 1984) are rimmed by distinctive Fe and Fe-Mn-rich oxide sediments, known respectively, as ochres (Robertson 1976) and umbers (Robertson & Hudson 1973). Ochres are strongly enriched in Fe and Cu, but depleted in A1 and other terrigenous-related elements (e.g. Ti). Umbers are rich in Fe and Mn and many trace elements (e.g. Cu, Ni, Pb, Co) relative to average deep-sea clays, but strongly depleted in terrigenous constituents. Both the umbers and ochres show negative Ce anomalies, reflecting rapid precipitation from seawater without selective fractionation of the rare earth elements (Fleet 1983). In addition, there is another less well known class of hydrothermal deposit that is strongly enriched in Mn and associated trace elements (e.g. Ba, Sr), but depleted in terrigenous constituents and related trace elements (e.g. A1, Zr). This type of hydrothermal deposit,

that is the subject of this paper, is strongly manganiferous, with very subordinate Fe and trace metal contents. Recent work suggests that most hydrothermal heat flux to the oceans takes place away from spreading axes, with less than 2.5% of the total being related to high-temperature hydrothermal activity on spreading ridges (Ginster et al. 1994; Bemis et al. 1993; Baker et al. 1993; Rona & Trivett 1992). Most of the remaining heat loss relates to low-temperature-type hydrothermal discharge, in the vicinity of spreading ridges and in various off-axis settings. Low-temperature hydrothermal systems are defined as those for which the temperature of the venting fluids is < 150°C, compared to high-temperature systems, with temperatures of up to 350-400°C (Edmond et al. 1979; German et al. 1995; Mills 1995). Low-temperature hydrothermal fluids commonly precipitate Mn-fractionated oxide sediments in the vicinity of source vents, commonly as crusts (Scott et al. 1974; Shearme

ROBERTSON,A. & DEGNAN,P. 1998. Significance of modern and ancient oceanic Mn-rich hydrothermal sediments, exemplified by Jurassic Mn-cherts from Southern Greece. In: Mn, Ls, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 217-240

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et al. 1983; Rona et al. 1984; Thompson et al.

1985). Where high-temperature hydrothermal activity occurs at depth beneath the seafloor (within lavas or sediments) Fe can be extracted from the circulating fluids to form sulphides, leaving a lower-temperature effluent rich in Mn that is then exhaled onto the seafloor and precipitated as Mn-rich oxyhydroxides (Scott et al. 1974; Bonatti et al. 1972). In addition, there are hydrogenetic (synonymous with hydrogenous) deposits (e.g. Mn nodules) which are typically relatively enriched and Mn, Fe and many trace elements (e.g. Co) (e.g. Bonatti & Nayudu 1965; Price & Calvert 1970; Glasby 1977). Fe-rich ochres precipitate in the vicinity of high-temperature black smokers (i.e. metres to tens of metres away), whereas Fe-Mn rich umbers accumulate from fall-out of metalliferous particles at greater distances (i.e. hundreds of metres to several kilometres away). The latter type of dispersed sediments take the form of A1poor, Fe-Mn-rich and trace-metal enriched sediments that are distributed over large areas of the flanks of spreading ridges, notably the East Pacific Rise (Bostr6m & Peterson 1966; Bostr6m 1973; Bonatti 1975; Barrett et al. 1988). Fe-Mn-rich metalliferous umbers also overlie oceanic crust drilled by the DSDP (Dymond et al. 1973; Cronan 1976a). Hydrothermal effluent may also be transported long distances by currents, as neutrally buoyant plumes, before settling as dispersed manganiferous precipitates, hundreds to several thousand kilometres away (Lupton & Craig 1981: Trefry et al. 1985; Klinkhammer et al. 1985). The presence of a high-temperature-derived hydrothermal component is revealed, particularly by relative enrichment in Fe and Cu, whereas low-temperature-type deposits are generally enriched in Mn, but depleted in trace metals. Fe is separated from Mn in hydrothermal plumes, as Fe is more rapidly oxidized and precipitated than Mn (Krauskopf 1957; Bignell et al. 1976; Cronan 1976b; Edmond et al. 1979). This process can give rise to Mn enrichment far from high-temperature-type source vents. In addition, where widely dispersed, precipitates from both high- and low-temperature-type effluents become enriched in diagnostic trace metals (e.g. Co) derived from seawater. Using solely chemical evidence, hydrothermal deposits derived from low-temperature-type Mn exhalations are not always easily distinguishable from widely disseminated Mn-rich precipitates from high-temperature-type vents, especially where hydrothermal constituents are widely dispersed through pelagic or hemipelagic sediment. Where such uncertainty exists, field

relations provide an additional indication of hydrothermal type. Low-temperature-type hydrothermal activity is generally of low intensity and may be diffuse compared to black smoker exhalations (Scott et al. 1974; Baker et al. 1993; Bemis et al. 1993). Thus, resulting Mn-oxide precipitates tend to be located near source vents (i.e. metres to several kilometres away). By contrast, far-travelled precipitates from hightemperature vents are generally dispersed within pelagic or hemipelagic sediment away from igneous basement. This paper has three related objectives. The first is to emphasize the range and diversity of Mn-rich hydrothermal sediments in the oceans. The second objective is to examine a case history of Mn-rich cherts (i.e. lithified siliceous sediments) from the sutured Mesozoic Tethys ocean (Neotethys) in southern Greece, based on field and chemical data. The third objective is to discuss the importance of Mn-rich hydrothermal sediments in other areas of the Mesozoic Neotethys and elsewhere.

Oceanic Mn-rich deposits We begin by outlining some of the main occurrences of Mn-rich hydrothermal sediments in the oceans, including ridge axis, ridge flank and back-arc settings. In the TAG hydrothermal field (26°N, Mid Atlantic Ridge), low-temperature-type Mn precipitation takes place in the axial zones, within 5 km of an area characterized by high-temperature hydrothermal activity (Scott et aL 1974; Rona et aL 1984; Thompson et al. 1985; Mills 1995). These deposits contain very high abundances of Mn (>30%), but low Fe (<0.15%) and trace metals (Co < 50 ppm, Cu < 120 ppm), and are inferred to have precipitated rapidly, mainly as localized crusts (Scott et al. 1974). According to Rona et al. (1984) fractionated Mnoxide trace element-poor sediments cover about 10% of the seafloor within the TAG hydrothermal field and are also distributed along faults sub-parallel to the rift valley. By contrast, Fe-oxide sediments (ochres) and sulphides in the TAG area are strongly depleted in Mn relative to the Mn-deposits (Shearme et al. 1983; Metz et al. 1988; Klinkhammer et al. 1986). Mn-rich oxide sediments are known to precipitate away from the spreading axis (tens to hundreds of kilometres away), notably in the Galapagos area (Londsdale 1977; Moore & Vogt 1976; Malahoff et al. 1983) and the Juan de Fuca Ridge (Bemis et al. 1993). Off-axis hydrothermal fluids may interact with pelagic or hemipelagic sediments

MODERN AND ANCIENT OCEANIC MN-RICH HYDROTHERMAL SEDIMENTS

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overlying the oceanic basement, giving rise, for example, to the Galapagos h y d r o t h e r m a l mounds (Corliss et al. 1978, 1979; Crerar et al. 1982; Honnorez et al. 1983). Many of the above deposits are of low-temperature type and precipitated close to source vents. Mn-rich hydrothermal oxide-sediments are also found in a number of active arc and backarc basin settings, including the SW and W Pacific (e.g. Cronan et al. 1982; Moorby et al. 1984; Usui et al. 1986; Usui & Someya 1997) and the Mediterranean (e.g. Tyrrhenian Sea; Rossi et al. 1980). A good example is the Eastern Manus Basin (W Pacific) where high-temperature vent material is locally coated with Mn oxides, of inferred low-temperature origin (Scott & Binns 1995). Ferromanganese oxides sampled from the Valu Fa, Tonga and Lau Ridges (Hein et al. 1990) include deposits in which Mn is strongly fractionated from iron. Actively growing manganese-rich crusts from the Tonga islands receive Mn from the Tofua volcanic arc and the adjacent Valu Fa spreading centre. The crusts grow rapidly and are inferred to be < 1 Ma old. Diagenetic modification has taken place during growth. In addition, there are a number of examples of hydrothermal Mndeposits related to now inactive arc systems in the SW and W Pacific (Usui et al. 1986; Usui & Someya, 1997). Hydrothermal-enriched sediments of the Lau

basin, SW Pacific contain up to 11.6% Mn and up to 9.82% Fe (Hodkinson & Cronan 1995). The Fe-Mn oxides are interpreted as distal fallout from high-temperature axial vents, although local metal anomalies could relate to off-axis activity. These metalliferous sediments are slightly enriched in Mn relative to Fe compared to average East Pacific Rise hydrothermal sediments. Ferruginous oxide-sediments (depleted in Mn) are also reported from a number of off-axis seamount-type settings, both in the open oceans and in marginal basins. Deposits associated with mid-Pacific seamounts are dominated by iron oxyhydroxides (mainly goethite), with relative depletion in Mn, low abundances of trace elements and REE. Iron-rich hydrothermal sediments also occur on the Dellwood Seamount, N W Pacific, the Loihi Seamount (Hawaii), on several seamounts in the S Pacific and on seamounts associated with the East Pacific Rise spreading centre (Hein et al. 1994). These deposits are interpreted as proximal (near source) low-temperature-type deposits formed by chemical interaction of hydrothermal effluent with pelagic sediment beneath the seafloor (Hein et al. 1994). To explain the paucity of Mn, it is probable that any hydrothermal Mn was dispersed away from the seamounts. Strong Mn fractionation is apparent in many settings in the modern oceans. Discoveries

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appear to be mainly in rift and back-arc basins, but this may be because exploration of spreading ridges has focused on high-temperature-type black-smoker fields at the expense of peripheral and off-axis areas. Mn-rich, Fe-poor primary hydrothermal deposits form from low-temperature hydrothermal activity both near spreading axes (e.g. TAG), and in off-axis settings (e.g. Tonga).

Case history: Mn-cherts from the Neotethys ocean in Southern Greece

Geological setting Having highlighted the importance of Mn-rich hydrothermal deposits in the oceans, we now consider evidence that comparable Mn-rich deposits are present in the stratigraphical record of ancient oceans preserved as fragments within orogenic belts on land. We take as an example Jurassic Mn-rich siliceous deposits (cherts) from the sutured Mesozoic Neotethys ocean in the Peloponnese area of southern Greece (Figs 1 & 2). These deposits are mainly Mn-rich cherts that occur together with tholeiitic volcanics and

deep-sea sediments. However, they are not associated with ophiolites, unlike other better known examples of Neotethyan metalliferous sediments (e.g. in Cyprus and Oman; Robertson & Boyle 1983).

Regional setting The Neotethys ocean formed by rifting of North Africa (Gondwana) in the Triassic and remained partly open until the Early Tertiary (e.g. Robertson et al. 1991). Deep-sea sediments accumulated on an inferred oceanic basement of Late Triassic and younger Mesozoic age, as seen in the Peloponnese area of southern Greece (Pindos zone, Fig. 3). The igneous basement was subducted during closure of the Neotethyan oceanic basin in the Early Tertiary, leaving only small fragments in a volcanic-sedimentary melange (Degnan & Robertson 1998a), whereas the overlying sediments were accreted as substantial thrust sheets within an accretionary prism that was finally tectonically emplaced over the adjacent continental margin (Adria). These thrust sheets comprise long-ranging successions of Late Triassic to Early Tertiary age

MODERN AND ANCIENT OCEANIC MN-RICH HYDROTHERMAL SEDIMENTS

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(Fleury 1980; Baltuck 1982; De Wever 1989) and include Mn-rich siliceous deposits (cherts) (Degnan & Robertson 1998b). Complete sedimentary successions within individual thrust sheets include radiolarian cherts, argillites, redeposited carbonates and turbiditic quartzose sandstones, together, ranging in age from Late Triassic to Early Tertiary. Tens of imbricate thrust sheets are present, with the Mn-cherts discussed here being located at various structural positions within the overall stack of thrust sheets (Fig. 4). The Mn-cherts occur within a succession of radiolarian cherts and argillites, depositionally overlying tholeiitic volcanic rocks at several localities S E of Patras (Degnan 1992; Fig. 1). At one locality (Kombigidi, Fig. 3), cherts interbedded with these lavas are dated as Late Jurassic using radiolarians (Degnan 1992). The manganiferous cherts are associated with regionally extensive radiolarian-rich cherts of Middle-Late Jurassic (Bajocian - Tithonian) age (Degnan and Robertson 1998b). The cherts are interpreted to relate to regional highproductivity within Neotethys during Mid-Late

Jurassic time (Fleury 1980). The Late Jurassic volcanic rocks and associated Mn-cherts can be restored, using field structural and stratigraphical criteria, to a location on the palaeo-abyssal plain of the Neotethys (Pindos) ocean between a passive margin and a spreading axis. This restoration rests on the assumption that Tertiary tectonic emplacement took place by simple piggy-back thrusting (in-sequence thrusting). The most proximal (near margin) units are at the base of the thrust stack, whereas the most distal (more axial) units are at the top of the thrust stack (Degnan 1992; Fig. 4). Thrust sheets, both above and below Jurassic volcanics and Mn-cherts include successions that extend downward into a Late Triassic interval and upward into a Cretaceous-Early Tertiary succession (Fig. 4). The significance of this observation is that the Mn-cherts accumulated in a deep-sea basin that originated prior to the time of Late Jurassic volcanism and persisted until early Tertiary time. Regional studies indicate that the Jurassic volcanics were erupted onto the floor of an

222

A. ROBERTSON & P. DEGNAN Central

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Fig. 4. Structural cross-section across part of the thrust belt showing the approximate positions of the metalliferous sediment localities discussed here. Drimos and Aroania lie quite close to this line of section, but Kombigadi is located several kilometres further north (See Fig. 3).

oceanic basin that rifted in the Early-Mid Triassic, followed by sea-floor spreading in Late Triassic-Jurassic time (Robertson et al. 1991). The Jurassic volcanics and Mn-cherts discussed below are thus seen as having formed in an offridge-axis setting within an oceanic basin. The width of the Neotethyan ocean basin in which the Mn-cherts accumulated is unknown, but is estimated to be not more than 1000kin. This particular oceanic basin (Pindos ocean) was bordered by continental fragments, of comparable palaeogeographical complexity, for example, to the modern S W Pacific region. After deposition in the Late Jurassic the succession was subjected to burial by several hundred metres of deep-sea sediments of CretaceousEocene age, followed by tectonic deformation within a thrust stack several kilometres in structural thickness (Degnan 1992). Thrust faulting was followed by regional N-S orientated extensional faulting, at right angles to the direction of opening of the Gulf of Corinth to the north. Most recently, exposure and weathering has taken place in a Mediterranean mountainous climate.

M a n g a n e s e enrichment At a locality where the lavas and associated volcaniclastic sediments are dated as Late Jurassic (Kombigadi, Fig. 3) they are depositionally overlain by argillaceous radiolarian cherts, but with only minor manganese enrichment. These lavas occur at the original base of a chert succession (possibly locally overturned) by thrusting. At a second locality, Aroania, located in a higher thrust sheet (Fig. 3), similar lavas and

volcaniclastic sediments are depositionally overlain by Mn-cherts, dated as Late Jurassic using radiolarians (G. Rose in Degnan 1992). This occurrence is as a small isolated thrust sheet in which no overlying succession is exposed. At a third locality, Drimos, located 4 km away, in a slightly higher thrust sheet (Fig. 3), numerous thin layers (5-20 mm) of Mn-chert are present within an undeformed chert succession. Intact stratigraphically overlying successions are preserved in the vicinity. The local successions and inferred composite successions are shown in Fig. 5. In the past, ore-grade horizons were mined at Drimos and Aroania (Galanopolous 1982). Elsewhere, layered Mn-rich cherts of less than ore grade alternate with argillaceous cherts within intact stratigraphical successions at numerous other locations throughout the N E Peloponnese (e.g. Lambia, Agridi, Lesteena, Figalia, Kontovazani, Platanitza, Drakovouni, Likouria; Fig. 3), but there is no evidence of locally preserved depositional contacts between cherts and volcanic rocks at these localities (Degnan 1992; Degnan & Robertson 1998b). Here, we focus on the best exposed locality, Aroania, where the Mn-cherts are spatially associated with volcanics (Fig. 6). The base of the succession, as exposed, is volcaniclastic breccia, including sub-rounded blocks of basalt and lapilli (up to 1 m in size). This material is well cemented by several generations of silica. Whole-rock geochemical analysis of major and trace elements shows that the lavas are of basic to intermediate composition and belong to a variably fractionated tholeiitic suite that was affected by hydrothermal alteration and weathering (Degnan 1992).

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Fig. 5. Stratigraphic setting of the Mn-chert in S Greece. (a) Complete succession, where preserved. The Mndeposist are associated with the Late Jurassic interval. (b) Local successions on Mn-cherts: (i) Kombigadi, (ii) Aroania, (iii) Drimos (Not to scale). See Fig. 3 for locations. At Aroania, volcaniclastic sediment is depositionally overlain by thinly-bedded argillaceous cherts, ranging in colour from dark red to brown, orange and pink. These cherts are interbedded with numerous layers of dark grey to black Mn-chert. Although sheared, these cherts are observed to form thin beds up to several centimetres thick, together with millimetre-scale laminae within individual thin beds (Fig. 7) that can be traced up to several metres laterally. The Mn-chert also contains minor amounts of volcaniclastic siltstone and sandstone. In addition, there is a chert breccia 12 cm thick (Fig. 6) that contains small, angular clasts of dark Mn-rich chert in an orange clay-rich, Mn-poor matrix. Syn-depositional faulting, dewatering structures and small-scale slumping indicate contemporaneous tectonic instability. Fine-scale Mn-rich and Mn-poor alternations precisely follow depositional laminae. This observation and the presence of the syn-depositional chert breccia strongly suggest that the manganese enrichment was syn-depositional rather than the result of diagenetic enrichment.

Results and discussion

Petrography and mineralogy Where concentrations are highest, the manganese has a black speckly lustre. In thin section, the Mn-chert is dominated by finely crystalline chalcedonic and cryptocrystalline quartz with zones of opaque manganese oxide. The only manganese mineral definitely identified by whole-rock X-ray diffraction is pyrolusite. However, delta MnO2, possibly representing disordered birnessite, was identified by the occurrence on X-ray diffraction traces of weak peaks at 1.42A and 2.4_A (Burns & Burns 1977). Haematite and goethite are also locally present. Other minerals identified are illite, calcite and apatite.

Whole-rock geochemistry Twenty-eight samples of Mn-cherts were analysed by X-ray fluorescence for major-element oxides and trace elements. Absolute abundances

224

A. ROBERTSON & P. DEGNAN

Fig. 7. Photomicrograph of hand specimen showing alternating Mn-rich and Mn-poor laminations. The lamination is inferred to be of primary origin. Part of this sample was analysed using backscattered electron microscopy (see Fig. 11). Scale bar 2 cm.

and ratios of Fe203/A1203 and Fe203/MnO are given in Table 1. Samples are mainly from the ore-grade deposits (Aroania and Drimos), but samples of Mn-rich sediments from other localities are included. Most of these are cherts, but some contain lower levels of silica than defined for chert (owing to high contents of MnO) and are strictly Mn-oxide sedimentary rocks. However, for brevity, the general term Mn-chert is used here for all the sediments analysed. The most important features of the chemical data are as follows:

J

CSFMCGC KEY Mn-rich chert Siliceous mudstone Mn-chert syn-sedimentary breccia Siticified volcaniclastic sediment Mafic/intermediate extrusive rock • Shear planes

Fig. 6. Detailed sedimentary log of the Mn-chert occurrence at Aroania , the most important of the three Mn-chert localities discussed. See Fig. 3 for location.

(i) MnO ranges from 2.4-69.9%, whereas levels of Fe203 are generally lower, ranging up to 4.95%; (ii) A1203 values are low and range from 0.32.4O%; (iii) Trace elements are in the following ranges: Ba (10-1199 ppm), Ce (up to 58 ppm), Ni (1-270 ppm), Cu (11-2079 ppm) Zn (5-87 ppm) and Pb (up to 197 ppm); (iv) Normalised abundances of selected trace elements are within, or very close to, the range of values known for oceanic hydrothermal and combined hydrothermal and diagenetic occurrences (Bonatti et al. 1972); (v) Ratios of FezO3/A1203 and Fe203/MnO show marked variation (Table 1). At Aroania, both Mn-rich ( ~ 2 4 % ) and relatively Mn-poor cherts (3%), are present, whereas at Drimos all but one of the samples analysed is highly Mn-rich. Samples from other localities include both more and less Mn-rich sediments. Two samples from Aroania (Fig. 3),

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33 2897 11 24 3 0.43966 0.17508

9 35 9 110 20 30 15

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90.53 1.16 0.24 0.80 0.80 0.11 0.26 0.05 2.91 0.03 3.21

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Table 1. Major- and trace-element compositions of mangan(ferous cherts and manganiferous oxide-sediments. Locations of rocks analysed. 63 Drimos, 64 Drimos, 65 Drirnos, 66 Drimos, 67 Drimos, 68 Drimos, 69 Drimos, 80 Aroania, 81 Aroania, 90 Aroania, 94 Aroania, 95 Aroania, 96 Aroania, 124 Lambia, 158 Aroania, 174 Agridi, 345 Lesteena, 347 Lesteena, 354 Figalia, 439 Kontovazani, 440 Kontovazani, 452 Platanitza, 486 Drakovouni, 504 Aroania, 510 Aroania, 600 Likouria, 601 Likouria, 614 Elatophyton, 702 Alepohorio. See Fig. 2 f o r locations. Method of analysis; see Fitton & Dunlop (1985)

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shales). However, hydrothermal-type sediments show very strong compositional variations (e.g. A1, Fe, Mn, Cu). Interpretations of different modes of formation of such sediments (e.g. hydrothermal versus hydrogenetic; see below) are based on strongly differing compositions and comparisons of the ratios of a range of constituents.

Electron microprobe and S E M study One sample of Mn-chert from Aroania (Fig. 7) was studied using electron microprobe in an attempt to identify the distribution of major constituents in relation to primary sedimentary lamination. The sample consists of millimetrescale laminations of black Mn-rich argillaceous chert, intercalated with orange argillaceous chert devoid of any significant Mn content. Twenty six analyses were made to establish the elemental variation across laminae and to identify detrital particles (Degnan 1992; Table 2). The elemental distribution is shown in a back-scattered electron image (Fig. 10). The area studied includes

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•

1

60

-.

•fi- -

20

MnO

40 MnO

A

q

60

E 300.

, •

,

2500 "-

l

250

2000

.-- 200 1500 Z 15o~/ 1oo ~-•

• • I OO0

500 .

°~ °

?o

20

000

~o

20

o

MnO

G

40 MnO

H

200

500 400

150

300 >

IOO

•

2OO 5O

1ooj D

•

•

•

. .

oRb.t/" o

2o =

•

;o

•

6o

~o

~'o

MnO

I

6'0

MnO

J

Fig. 8. Variation diagrams of major-and trace-elements. (E) Ba versus MnO; (F) Sr versus MnO; (G) Ni versus MnO; (H) Cu versus MnO; (I) Zn versus MnO; (J), V versus MnO. Axes are % of total composition for oxides and ppm for trace elments.

part of a dark Mn-enriched lamina (upper, layer A), a transitional zone between a Mn-rich and a Mn-poor lamina (layer B) and part of a Mnpoor lamina (layer C). Most of the points analysed are shown in Fig. l la. The relative abundances of SiO2, M n and A1203 for each point analysed are shown in Fig. 1 lb. The results confirm that the M n is concentrated in primary laminae, rather than being patchily distributed throughout the rock during diagenesis.

In addition, a Mn-rich lamina and transitional zone were studied using scanning electron microscopy to reveal the morphology of the mineral phases present (Fig. 12 a, b). The Mnrich zone is dominated by bladed spherules of pyrolusite, with tiny nuclei mostly composed of silica. The transitional zone is dominated by approximately spherical mineral bodies that exhibit an internally zoned colloform texture and a bladed mineral fringe (too fine to analyse).

0.356 0.124 0.544 58.334 0.167 2.125 0.27 28.031 0.524 0.014

90.49

Total

55.504

0.057 0.681 4.544 6.204 0.076 43.777 0.02 0.085 0.054 0.005

504/15

504/14

Na20 MgO A1203 SiO2 K20 CaO TiO2 MnO FeO NiO

96.516

96.632

Total

0.162 0.405 2.74 90.308 0.761 0.115 0.272 0.212 1.513 0.029

504/2

0.034 0.015 0.144 96.302 0.028 0.027 0.012 0 0.042 0.029

504/1

NazO MgO A1203 SiO2 K20 CaO TiO2 MnO FeO NiO

Sample PD/89/504

97.735

0.225 0.118 0.595 85.471 0.129 0.708 0.122 10.068 0.282 0.018

504/16

94.558

0.104 1.995 4.47 83.618 1.225 0.141 0.045 0.156 2.759 0.043

504/3

93.325

0.278 0.088 0.316 66.234 0.035 1.714 0.197 24.005 0.43 0.029

504/17

79.109

0.624 0.274 1.374 14.669 0.208 4.157 0.771 55.614 1.365 0.053

504/4

77.796

0.716 0.245 0.903 11.447 0.193 4.228 0.595 58.309 1.149 0.01

504/18

100.784

0.14 0.071 0.287 93.5 0.043 0.424 0.12 6.02 0.151 0.028

505/5

78.475

0.515 0.327 2.906 10.322 0.043 4.391 0.801 58.221 0.906 0.045

504/19

98.259

0.159 0.065 0.161 87.007 0.048 0.723 0.182 9.725 0.188 0.001

504/6

97.018

0.137 1.078 8.797 81.575 2.572 0.154 0.108 0.26 2.285 0.052

504/20

97.796

0.295 0.098 0.531 78.279 0.086 1.223 0.274 16.705 0.279 0.028

504/7

95.559

0.I82 0.066 0.614 87.967 0.082 0.47 0.14 5.867 0.17 0

504/21

97.487

0.217 0.091 0.372 84.172 0.083 0.86 0.198 11.221 0.247 0.024

504/8

93.198

0.332 0.09 0.435 68.205 0.089 1.559 0.16 21.873 0.436 0.02

504/22

97.017

0.278 0.129 0.642 80.711 0.123 1.028 0.239 13.566 0.275 0.025

504/9

98.708

0.082 0.06 0.962 96.422 0.234 0.084 0.025 0.733 0.094 0.013

504/23

74.284

0.717 0.166 0.414 5.941 0.039 4.34 0.534 61.084 0.979 0.071

504/10

81.822

0.042 0.09 1.298 0.965 0 0.434 0.02 78.736 0.223 0.015

504/24

101.224

0.158 1.766 17.81 73.335 4.858 0.168 0.123 0.457 2.506 0.043

504/11

91.321

9.328 0.133 12.884 68.267 0.278 0.083 0.033 0.035 0.271 0.009

504/25

52.925

0.105 0.098 0.896 31.353 0.205 1.494 0.185 18.152 0.418 0.019

504/12

89.904

0.276 0.121 0.45 58.905 0.687 1.599 0.202 26.988 0.659 0.018

504/26

69.558

0.257 0.083 0.697 53.735 0.113 1.119 0.145 13.115 0.268 0.025

504/13

Table 2. Electron microprobe analyses" o f Mn-rich chert sample from the Metalliferous sulphide- and oxide-sediments associated with spreading centres in the modern oceans can now be closely related, in terms of chemical composition and depositional setting, to counterparts in ophiolites

MODERN AND ANCIENT OCEANIC MN-RICH HYDROTHERMAL SEDIMENTS Si

Mn

AI

AI

Mn

K

Fe

L Ni+Cu+Zn x l O 0

Fe

229

Cu

Mn

M

Zr

Ni

N

Fig. 9. Geochemical plots continued. Ternary plots (K-M) highlighting the relative abundance of selected major and trace elements. In M, field A is hydrogenetic deposits, field B hydrothermal plus diagenetic deposits and field C is exclusively hydrothermal deposits. From Bonatti et al. (1972). The spheres are two orders of magnitude smaller than radiolarian tests. The analyses indicate that the spheres have a SiO2-rich core with some delta MnO and become progressively more MnO-rich outward, with a corresponding decrease in SiO~. The MnO increase is accompanied by small, but systematic, increases in CaO, TiO2 and FeO. Scanning electron and backscattered electron microprobe studies of modern manganese oxide crusts from the Mariana and Izu-Bonin (Schulz & Hein 1991) and West Antilles volcanic arcs (Kang & Kosakevitch 1984) reveal a range of textural features that include colloform texture, homogeneous dense layers, laminated dense layers, porous, fibrous and scaly morphologies. Schulz & Hein (1991) interpret amorphous cryptocrystalline manganese oxide as an initial phase of outward precipitation from amorphous cryptocrystalline cores or layers (commonly colloform). This morphology possibly represents periods of oxidation of Mn-rich fluids and rapid precipitation. Slower reaction kinetics favour slower growth of microcrystalline and crystalline morphologies. Alternations of cryptocrystalline and more crystalline-oxides might thus reflect episodic growth, as redox and other conditions varied. Modern comparisons thus support a primary origin for the Mn-rich laminae.

Source o f M n in chert Based on stratigraphical and structural evidence, the Mn-cherts in Southern Greece originally accumulated on an abyssal plain between a passive continental margin (Adria) and a spreading axis (Degnan & Robertson 1998b). During the Late Jurassic, tholeiitic volcanic rocks were erupted on this abyssal plain. The restoration suggests that all of the Mn-rich localities studied were originally < 100 km apart. The Mn-rich, thin beds and laminae within the argillaceous cherts, a priori, could have one of four origins: (a) detrital (terrigenous); (b) hydrogenetic (precipitates from seawater); (c) diagenetic (i.e. by secondary enrichment of Mn), or (d) Hydrothermal (either proximal or distal). Terrigenous origin. The Mn cannot mainly be present as detrital particles of terrigenous origin because: (i) There is no physical association with terrigenous material, as confirmed by optical and electron microscopy; (ii) A1203 and TiO2, both of mainly terrigenous origin, do not correlate with MnO abundances; (iii) Fe203 abundances are typical of average

230

A. ROBERTSON & P. DEGNAN

Fig. 10. Black and white print of false colour back-scattered electron image of laminated chert (sample in Fig. 7). The area studied is divided into three zones. (A) Part of a Mn-rich lamination; (B) The transition between a Mnrich and a Mn-poor lamination; (C) Part of a Mn-poor lamination. The large round shapes are radiolarians. Scale bar = 0.2 mm.

pelagic clay (Turekian & Wedhepohl 1961), but do not correlate with MnO in the Greek Mn-cherts. origin. Hydrogenetic processes include direct precipitation of metal oxides and adsorption of metal cations from seawater. Hydrogenetic open-ocean Mn nodules and crusts are typically strongly enriched in Ni, Co, Cu and REE, and exhibit strong positive Ce anomalies (Cronan 1976b; Glasby 1977; Fleet 1983; Hein et al. 1987b). The Greek Mn-cherts plot mainly in the hydrothermal plus diagenetic field, with some of the more Fe-rich samples plotting exclusively in the hydrothermal field (Bonatti et al. 1972; Fig. 9). Some Fe-Mn crusts (e.g. associated with the Tonga and Mariana arcs) are compositionally intermediate between hydrothermal and hydrogenetic deposits (Bon a t t i & Nayudu 1965; Hein et al. 1987b, 1990), and are interpreted as mixtures of hydrogenetic and hydrothermal constituents. The most Mndepleted (but more Fe-rich) Greek samples plot near the edge of the hydrogenetic field, suggesting that minor hydrogenetic enrichment has indeed taken place in these cases. Hydrogenetic

Diagenetic enrichment. Diagenetic mobilization

of Mn and related elements (e.g. Sr, Ba) plays a role in the formation of many manganese deposits, including manganese nodules and

crusts (Bonatti et al. 1972). Remobilization is most marked where reducing conditions develop upon burial, favouring dissolution of manganese and migration to an oxidizing interface, usually near the sediment surface (Price & Calvert 1970; Lynn & Bonatti 1965). However, several factors oppose large-scale diagenetic mobilization of manganese to form the Greek Mn-cherts considered in this study: (i) the Mn is concentrated in laterally continuous thin beds and laminations that are concordant with primary lamination, as defined by fine detrital grains (e.g. clay minerals). The manganese is not nodular, or patchily distributed as in typical diagenetic Mn deposits (e.g. Mn nodules); (ii) the interbedded argillaceous sediments are not mineralogically or chemically altered (e.g. by hydrothermal fluids from the oceanic basement); (iii) the more Fe-rich siliceous sediments plot outside the diagenetic field, defined by known occurrence of diagenetic manganese deposits (Bonatti et al. 1972); (iv) although containing Radiolaria, the associated argillaceous cherts are red or brown and are interpreted as oxidizing sediments (De Wever 1989) in which Mn, should not be readily mobile.

M O D E R N A N D A N C I E N T OCEANIC M N - R I C H H Y D R O T H E R M A L SEDIMENTS i

i oO __,,'2

0 0

0

Oo

t.,.-'

0 ,-,

o oO

o0o

0

0

0

0

0

0 o

~

f.;;:'.(51

o

0

0

O~ ~

0

5o4"~ou

o O~O

0504/H@

Pyrolusite

"~0

~U

o

0

o

°oo

c~

oo

'~";

Mn02

"*

-o.

Otz

o

m

° S~

504;4~

IIIite & Microcrystalllne

£

o

%i o Authigenic Otz &

o ~)0

0

c,

~ - o

0 v

o

~ o¢

o 0°.

O0 0

(.~ o o

231

o

504/2~ • 5{)4/3

----

504/1

o

o

SiO2 --

A

I 0.1ram

Si02

AI203

MnO

~i ~14

lo

13

I

1~4

lO

is

[-

14

~,o / k 13

11

(18%)

A12

q

g

~ 5

7

0.1 ~

4

23

'¢ I <

B

0

I

40

i

60

I

8 %

/X4 ~

mm

3

1

I

20

40%

I 1

I 2

I 3

I

i> 5%

> Laminae

1~504/1

Fig. 11. Map of a back-scattered electron image of laminated manganiferous chert. (a) showing the electron probe analysis points from part of a Mn-poor lamination, across a transitional zone and into a Mn-rich lamination (Zones A, B and C in Fig. 10). The minerals identified are marked on the left; (b) The analysed concentration of Si, Mn, and A1 oxides through the transect. Note how SiO2 and MnO vary in inverse proportion, whereas Al203 concentrations (largely silty clay) remain essentially constant. Very high SiO2 values (point 502/10) are observed where a radiolarian test was analysed. The main minerals present are shown on the left-hand side of Fig. 1 la. Note the enrichment in pyrolusite in the Mn-enriched zone relative to illite, and of microcrystalline quartz in the Mn-poor zone. Microcrystalline quartz and pyrolusite occur patchily in the transitional zone.

232

A. ROBERTSON & P. DEGNAN

A

B

Fig. 12. Scanning electron micrographs of (a) of Mn crystallization in the transitional zone of Fig. 10. The subspherical colloform mineralization consists of alternations of SiO2 and disordered MnO with a thin fringe of pyrolusite crystals; (b) Radiating acicular crystals of pyrolusite from part of Mn-rich lamination. The cores of the radiating bodies contain small amounts of SiO2. Both scale bars = 50 #m.

Hydrothermal origin. A mainly hydrothermal origin is favoured by the chemical evidence, as follows:

(i) strong enrichment of Mn relative to Fe203 and A1203; (ii) abundances of trace elements (e.g. Ce, Ni, Zn, Pb and Cu) are within the range of oceanic hydrothermal deposits (Bostr6m 1973); (iii) levels of Co in Jurassic Mn-cherts are within the range typical of hydrothermal deposits (i.e. < 300 ppm) for samples from a number of localities within the Pindos zone of S Greece (Varnavas & Panagos 1982; Panagos & Varnavas 1984). In contrast, hydrogenetic deposits are relatively enriched in Co (e.g. Schulz & Hein 1991). However, samples from Aroania and Drimos are enriched in Co relative to those from other localities in Greece (S. P. Varnavas, pers. comm. 1997). A small number of samples with relatively high values of Cu, Pb, Ni and Ce have undergone minor hydrogenetic enrichment, similar to some modern hydrothermal sediments (Hein et al. 1987b). The silica in the Greek Mn-cherts is assumed to be mainly biogenic, derived from radiolarians and other siliceous microfossils, in keeping with a setting within a radiolarian chert succession. This origin is supported by the microscopic evidence of preserved radiolarians in these sediments. The Mn-cherts are compositionally

similar to the composition of deposits from a range of hydrothermal environments. These include the strongly fractionated back-arc basin deposits in the S and SW Pacific described by Hein et al. (1990), in which absolute values of Mn reach ~45%, Fe < 1%, with A1 and trace element values (e.g. Cu, Ni, Pb) exhibiting similar values to the Greek Mn-cherts. Also, the Fe-Mn-Si deposits of the Woodlark Basin contain up to 2.7% MnO. 67.7% SiO2 and 2.93% total Fe203. Levels of A1203 and Ni in these deposits are similar to the Greek Mncherts, whereas Zn is slightly lower (Scott & Binns 1995). The silica in these sediments is inferred to be mainly of hydrothermal origin. By contrast with the above oceanic deposits, the Greek Mn-cherts are also highly enriched in silica that is assumed to be of biogenic origin related to high palaeoproductivity within Neotethys in the Late Jurassic (De Wever 1989).

Significance of ancient Mn-hydrothermal deposits In the above section we have interpreted the Greek Mn-cherts as Mn-rich hydrothermal precipitates within an area of high siliceous productivity. We now consider how important similar deposits are elsewhere in the sutured Neotethys and in other orogenic belts. The most directly relevant comparisons can be made with Mn-rich hydrothermal deposits else-

MODERN AND ANCIENT OCEANIC MN-R1CH HYDROTHERMAL SEDIMENTS EARLY JURASSIC

.,.,.-.

\

Fig. 13. Palaeo-oceanographical reconstruction of Neotethys in the Eastern Mediterranean region during Jurassic time; area of study shown in the box. Modified after Robertson et al. (1991).

where in Greece, as they formed within the same oceanic basin (Fig. 13). These, unlike the Mncherts described above are mainly spatially associated with ophiolites. Within the Argolis Peninsula of the eastern Peloponnese of S Greece (Fig. 2), east of the study area, MORBtype extrusives are depositionally overlain by thin (several metres) Mn-ores (Robertson et al. 1987) that are compositionally similar, in Fe/Mn ratios and trace metal content, to the Mn-cherts discussed above. A mechanism was proposed involving fractionation of Mn from Fe in a hydrothermal system dominated by relatively low-temperature venting. Elsewhere, in the Othris area of central Greece (Fig. 2), MORBtype ophiolitic extrusives are associated with small sulphide stockworks (Valsami 1990), massive sulphides, Fe-cherts (jaspers), Si-Ferich sediments and Fe-Mn umbers. The Fe-rich deposits (ochres) formed close to the massive sulphides (tens to hundreds of metres away), whereas Fe-Mn deposits (umbers) accumulated further away (hundreds of metres to several kilometres) (Robertson & Varnavas 1993a). Mnrich siliceous deposits (now cherts) also accumulated further from the sulphides (i.e. kilometres away). These Othris Mn-cherts exhibit

233

concentrations of Co (average 130 ppm) and other trace metals (Ni, Pb, Zn, Cu), similar to the Mn-cherts studied in S Greece (Panagos & Varnavas 1984), and are interpreted as hydrothermal deposits that accumulated away from the spreading centre (i.e. off axis). Similar massive sulphide and metalliferous oxide sediments occur within a continuation of the same ophiolite belt further N into Albania (Hoxha 1995; Bortolotti et al. 1996). Elsewhere in the Eastern Mediterranean area, Late Triassic metalliferous cherts and metal oxide-rich sedimentary rocks are spatially associated with basic extrusives within the Antalya Complex, S W Turkey (Robertson 1981). These lavas and hydrothermal deposits formed during the early stages of opening of the Neotethys ocean. The metal-rich sediments are interbedded with and overlie thick (up to 1000 m) basic lavas of both mid-ocean ridge (MORB) and withinplate type (WPB). These lavas represent oceanic crust formed in the initial stages of spreading to form a Neotethyan ocean basin. Plutonic ophiolitic rocks are assumed to exist at depth, but are not exposed. Mn-cherts of Jurassic-Early Cretaceous age are also found within tectonically emplaced deep-sea sedimentary successions in various areas of the Eastern Mediterranean, including the Antalya Complex, SW Turkey, the Mamonia Complex in SW Cyprus and the Baer-Bassit units of NW Syria (Robertson & Boyle 1983; Fig. 1). These sediments are similar in field setting and chemical composition to the Greek Mn-cherts and are not directly associated with ophiolites or other magmatic rocks. They comprise deep-water successions that accumulated on continental rise/abyssal plain settings adjacent to passive continental margins (Robertson 1981). This situation is similar to that of the S Greek Mn-cherts discussed earlier. Comparable Mn-rich hydrothermal deposits also occur in Neotethyan areas outside the Eastern Mediterranean region, notably associated with Mesozoic, emplaced deep-sea sediments beneath the Semail ophiolite. These include Mn-cherts and crusts of Late Triassic age, found overlying oceanic basalts, and interpreted as oceanic crust formed near a rifted continental margin (Robertson 1986). Alternatively, these deposits were associated with seamount volcanism in a more open-ocean setting. There are also occurrences of Jurassic Mn-rich radiolarian cherts which contain intercalations of Mn-chert. These cherts are located within thrust sheets of deep-sea sediments, interpreted to have accumulated on oceanic crust bordered by a passive continental margin

234

A. ROBERTSON & P. DEGNAN

(Arabian margin; Robertson & Searle 1990). Kickmaier & Peters (1991) proposed a model whereby Mn was continentally derived. The Mn was initially deposited in a proximal setting on the continental slope within the oxygen minimum zone. This setting favoured dissolution of Mn that was then remobilized by currents and carried to abyssal depths, scavenging trace elements en route. Hydrodynamic sorting during final deposition gave rise to fine lamination as exhibited by these Mn-cherts. Kickmaier & Peters (1991) discounted a mainly hydrothermai origin on the grounds that no spreading centre was located in the region during the Late Jurassic, prior to formation of the Semail Ophiolite in the Cretaceous. However, it is now widely accepted that spreading in Oman began in the Permian/Triassic and continued during the Jurassic (e.g. see Robertson & Searle 1990). As in Greece, the Jurassic sedimentary succession with the Mn-cherts was detached and accreted to form a thrust belt during emplacement onto a continental margin. By contrast, the oceanic basement was almost entirely subducted. There is, thus, no viable objection to the Oman Mn-cherts being interpreted as hydrothermal deposits, as in S Greece. Jurassic Mn-cherts are also well documented overlying the Apennine and Ligurian ophiolites in Italy. These ophiolites are interpreted as rifted oceanic crust of M O R type within the western Neotethys ocean (Barrett 1981). The Mn-cherts overlie basalt near the base of an overlying radiolarian chert succession. Layered Mn-deposits within chert are up to ~1 m thick and can be traced laterally for up to several hundred metres; they are concordant with bedding in the cherts down to a several-centimetre scale. In places, the layered Mn-cherts pass laterally into massive Mn deposits. F e - C ~ Z n sulphide deposits are present within the underlying volcanic succession, commonly close to Mn deposits (Bonatti et al. 1976). The Mn cherts are very strongly fractionated with Mn up to 46%, but very low in Fe (mostly < 1%). Levels of Co, Ni and Zn are low (tens of ppm), with Cu up to 950 ppm. Bonatti et al. (1976) suggest that high-temperature hydrothermal solutions reacted with the oceanic basement at depth, extracting Fe to form sulphides, whereas Mn was fractionated, emitted in solution and then precipitated on the surrounding ocean floor, explaining the presence of the Mn-deposits at the base of the sedimentary succession. The Apennine Mn-ores are very similar in chemical composition to the Mncherts from Drimos and Aroania in S Greece, although the burial history and resulting mineralogy differ.

Metalliferous metamorphic rocks

There are also numerous examples of Mn enrichment within metamorphic rocks of variable metamorphic grade and variable original tectonic setting. Some of these are potentially of oceanic hydrothermal origin. In N W Greece (Thessaly) amphibolites are spatially associated with zones of disseminated sulphide, small massive orebodies (metre-sized) and jaspers (Fe-cherts). The volcanics are overlain by thick marbles. The setting is interpreted as a Triassic? rift, with sulphides and proximal hydrothermal sediments, overlain by a shallowwater carbonate platform (Robertson & Varnavas 1993b). In the Cycladic islands, S Greece (e.g. Andros), quartz-rich pelitic schists are interbedded with three types of metalliferous metamorphic rocks, as described by Reinecke et al. (1985): (i) piedmont-spessartine and piedmont-rich quartzites with braunite lenses; (ii) spessartite quartzites; (iii) Na-rich pyroxene quartzites. Lithology (i) is enriched in Fe, Mn, Cu, Pb, Sr and As, relative to average terrigenous shale. The protolith of this lithology was interpreted by Reinecke et al. (1985) to be mainly terrigenous, with subordinate biogenic siliceous ooze, that was enriched in F e - M n possibly originating as Mn-nodules. The protoliths of (ii) and (iii) were possibly smectite-rich, affected by variable chemical segregation during metamorphism (Reinecke et al. 1985). However, we note here that the chemical composition of lithology (i) is more similar to hydrothermal F e - M n umbers than to manganese nodules. Discrete Mn-nodules are rarely observed in unmetamorphosed Neotethyan deep-sea siliceous sedimentary rocks. We interpret the schists as part of a subductionaccretion complex related to Early Tertiary subduction of Neotethys in the Greek area (Pindos ocean; Fig. 13). The absence of associated volcanic rocks with the metalliferous schists can be explained by subduction of an original underlying oceanic basement, whereas the overlying deep-sea sediments were preferentially accreted. We believe that these metalliferous metamorphic rocks originated as clay- and terrigenous-rich, siliceous, hydrothermal sediments on the abyssal plain of the Neotethys ocean, in a setting similar to their unmetamorphosed counterparts in S Greece. Outside the Neotethyan area, another wellknown example of metamorphosed metallifer-

MODERN AND ANCIENT OCEANIC MN-RICH HYDROTHERMAL SEDIMENTS Continental margin

Abyssal plain

235

Spreading ridge

A Seamounts

B

ca. 0 I

lOOkm I

West

East

Fig. 14. Alternative settings of hydrothermal Mn enrichment in the Greek Mn-cherts. (a) Derived from the spreading axis; (b) Derived from off-axis seamounts. Axially derived Mn-oxides contributed to pelagic sedimentation in the basin as a whole, but the preferred origin of the Mn-chert is as precipitates from off-axis (low-temperature?) hydrothermal vents.

ous cherts is found within the Franciscan Complex of Western USA (Koski et al. 1993). The Mn deposits there form ore lenses within radiolarian sediments that are rarely spatially associated with volcanics. The Franciscan Complex has partially undergone high-pressure metamorphism. The ores are Mn-rich but Fepoor and are dominated by rhodochrosite, manganese silicate, opal-CT and quartz. The orebodies are disc shaped, up to several hundred metres in diameter and tens of metres thick (e.g. Ladd east and Buckeye north bed). The deposits range from massive to interstratified with chert, secondary silica, or manganese silicates. Boundaries are sharp; adjacent cherts contain disseminated rhodochrosite, and layers of laminated silic~rhodochrosite-manganese silicates occur at the contact of some deposits. Massive chert beds adjacent to the Mn-ores are interpreted as forming when siliceous host rocks were partly replaced by carbonate. The average oxide ore contains 44.2% MnO, 1.7% Fe203 and 2.3% A1203, and levels of Cr, Cu, and Ni, similar to the Greek Mn-chert studied here, whereas Zn is much higher (c. 400 ppm). The origin of the Franciscan Mn-ores is controversial. Some authors explain the Mn enrichment in terms of diagenetic mobilization of Mn (associated with migration of hydrocarbons) from within organic-rich radiolarian sediments originally formed within a zone of nutrient upwelling near a continental margin (Hein et al. 1987a). The Mn was ultimately continentally derived a n d / o r volcanogenic. However, more recently these Franciscan sediments were re-interpreted as part of a tectonic

slice complex, in which many of the sub-parallel lithological contacts that were originally thought to be depositional are in fact thrusts (Blake et al. 1988). The high-pressure rocks of the Franciscan Complex are now widely interpreted as an accretionary prism related to subduction of oceanic crust and accretion of its sedimentary cover (Wakabayashi 1992). The paucity of igneous rocks associated with the Mn-ores, can again be explained by subduction of the oceanic basement as in Neotethyan examples discussed above (e.g. S Greece, Oman). Massive sulphides within the Franciscan Complex can then be attributed to a spreading ridge setting (Koski et al. 1993). The Mn-ores are similarly likely to record dominantly hydrothermal F e - M n input in an oceanic setting (Crerar et al. 1982; Chyi et al. 1984), followed by subduction/accretion. Setting o f hydrothermal M n discharge In the above extensive review we have shown that the palaeogeographic occurrence of the Greek Mn-cherts is not an isolated case, but that comparable examples of both spreading centre and off-axis-type Mn-rich hydrothermal sediments existed elsewhere in Neotethys and in other preserved oceanic areas. The main reason for a hydrothermal origin not being considered in the past is that the Mn-rich sedimentary successions are detached from former oceanic basement by subduction-accretion processes. In nearly all cases, the Mn-deposits cannot now be related to any preserved hydrothermal vents as the oceanic basement was subducted, in contrast, e.g. with the Fe-rich ochres and Fe-Mn-

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rich umbers associated with intact ophiolites (e.g. Troodos and Oman ophiolites). Can a far-travelled origin for Mn-rich hydrothermal sediment versus an origin in close proximity to a hydrothermal vent be inferred from geochemical evidence alone? Limited evidence from oceanic Mn-oxide deposits suggests that local deposits are relatively depleted in trace metals and show negative Ce anomalies, whereas dispersed sediments are relatively enriched in trace elements, including Cu, Co, Ni and Pb, transitional in composition to hydrogenetic deposits (Scott et al. 1974; Hein et al. 1987b). Local deposits in which sulphides precipitate beneath the surface may show enrichment in sulphide-related metals (e.g. Cu). Using these criteria, the Greek Mn-cherts represent relatively locally precipitated Mn-deposits, with minor hydrogenetic enrichment. To explain the origin of the Mn-chert in S Greece, two possibilities are: (1) The Mn-oxide-rich sediments drifted from the spreading axis and accumulated as a passive drape over deep-sea sediments and localized off-axis volcanics (Fig. 14a). In this case, the Mn was disseminated from axial high-temperature type vents (black smokers); or, alternatively, (2) The oxide-rich sediments relate genetically to off-axis volcanism (Fig. 14b). In this case, the Mn was probably derived from localized off-axis low-temperature-type vents. In a small ocean basin, as inferred for the Greek Neotethys (Pindos ocean), it may be impossible to discriminate, geochemically~ between these two settings, as compositionally similar Mn-cherts exist in both axial (e.g. Othris area) and off-axis settings (e.g. S Greece). However, any interpretation must explain why very Mn-enriched cherts occur only locally, in some cases spatially associated with contemporaneous volcanism. Elsewhere, throughout western Greece as a whole (i.e. the Pindos zone) ubiquitous Late Jurassic radiolarian cherts contain only minor enrichment in MnO (< 1%). The Mn is patchily distributed in these cherts and discordant with respect to individual layers, suggesting that diagenesis has played an important role in concentration of this Mn above typical abundances in deep-sea clays (Pe-Piper & Piper 1989). By contrast, the localized Mn-cherts studied here are enriched in MnO~ commonly by more than an order of magnitude relative to regionally occurring cherts. In summary, we propose the following interpretation for the Greek Mn-cherts: during the

Mid to Late Triassic and Jurassic, seafloor spreading took place in the Greek Neotethys (Pindos ocean), associated with formation of axial hydrothermal deposits. Volcanism then took place in the Late Jurassic away from the spreading axis (several hundred kilometres away?). Mn was precipitated from hydrothermal solutions emanating from vents associated with this volcanism. This hydrothermal activity was probably of low-temperature type. The Mncherts formed by relatively localized deposition (kilometres to several tens of kilometres away) in a setting of high siliceous productivity. Minor enrichment of hydrogenetic constituents took place where hydrothermal precipitates were dispersed from source vents. There are probably no major differences in the nature of Mn-rich hydrothermal sediments between the present-day oceans, for which only sparse records of Mn-oxide rich hydrothermal deposits exist, and counterparts preserved in orogenic belts (e.g. Neotethys), in which such Mn-deposits are quite common. However, such deposits are easily identified on land, compared to the deep ocean that remains largely unexplored for hydrothermal deposits except for rare local areas. In the oceans, exploration for hydrothermal vents has focused on mid-ocean ridges and other sediment-starved settings (e.g. back arcs) that have an inherently low preservation potential in orogenic belts. In contrast, the stratigraphic record is biased towards preservation of off-axis abyssal plain sediments that are preserved within accretionary prisms. One implication of the presence of oceanic hydrotherreal Mn deposits emplaced onto land is that significant discoveries of Mn-precipitating hydrothermal vents probably remain to be made in various off-axis oceanic settings.

Conclusions (1) Review of oceanic metalliferous sediments suggests that Mn-rich oxide- sediments represent an important type of hydrothermal deposit that has not been widely investigated. There are modern occurrences in a wide variety of settings including midocean ridges, off-ridge axis areas, back-arc basins and active arcs. (2) A case history, utilizing Jurassic Mn-cherts from the sutured Neotethyan ocean in S Greece, shows that Mn-rich cherts are locally associated with Late Jurassic volcanics and are strongly enriched in MnO, but exhibit only modest enrichment of trace metals (e.g. Cu, Ni, Pb) relative to deep sea-clays.

MODERN AND ANCIENT OCEANIC MN-RICH HYDROTHERMAL SEDIMENTS (3) Mn-rich hydrothermal deposits commonly occur in other Neotethyan units, including Greece (Argolis and Othris), Turkey (Antalya), Cyprus (Mamonia) and the Middle East (Oman). Such deposits are particularly associated with deep-sea sediments that accumulated on abyssal plain areas adjacent to a passive margin. (4) Where a hydrothermal origin has not generally been recognized for Mn-oxide deposits, in some cases this is because the original oceanic basement was removed by subduction, whereas the overlying sediments (including metalliferous sediments) are preserved in accretionary prisms (e.g. in Oman and California). (5) Some metamorphosed Mn-deposits can be interpreted as oceanic hydrothermal deposits of variable type and composition (e.g. Cyclades, S Greece). Typically these were accreted at subduction zones and later exhumed. (6) The S Greek Mn-cherts are interpreted as low-temperature h y d r o t h e r m a l deposits with minor hydrogenetic enrichment. The most probable origin is related to Late Jurassic off-axis volcanism within the Neotethyan ocean. (7) Apparent differences in the relative abundances of Mn-rich hydrothermal sediments in the oceans and preserved in orogenic belts are probably artefacts of exploration and preservation. Spreading axes with high-temperature-type hydrothermal deposits (massive sulphides and Fe-Mn umbers) are rarely preserved in orogenic belts owing to subduction. By contrast, off-axis Mn-rich low-temperature-type deposits within deep-sea sediment successions (little explored in the oceans) have a relatively high chance of final preservation within an accretionary prism. The first author's fieldwork was assisted by a grant from the Carnegie Trust for the Scottish Universities. The second author acknowledges funding by a NERC Studentship held at the Department of Geology and Geophysics of Edinburgh University. S. P. Varnavas (Patras University) is thanked for valuable discussion. We thank G. Angell, D. James and S. Kearns for assistance, respectively, with the X-ray diffraction, Xray fluorescence and microprobe studies. Helpful comments in the manuscript were received from two anonymous referees. References

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Ancient vent chimney structures in the Silurian massive sulphides of the Urals R. J. H E R R I N G T O N

1, V. V. M A S L E N N I K O V

2, B. S P I R O 3, V. V. Z A Y K O V 2 &

C. T. S. L I T T L E 1

1The Natural History Museum, London S W 7 5BD, UK 2The Institute of Mineralogy, Russian Academy of Sciences, Miass, Russia 3NIGL, Keyworth, Nottingham NG12 5GG, UK Abstract: Exceptionally well preserved volcanogenic massive sulphide ores at Yaman Kasy

in the Silurian of the southern Urals have yielded not only well-preserved sulphidized vent macrofauna but also fragments of vent chimneys. All fragments show a broad 3-fold mineralogical zonation. An outer zone which forms the chimney/conduit wall comprises largely pyrite and marcasite which is laminated or collomorphic and is commonly porous. In one fragment this zone is characterized by a honeycomb-like structure in the pyrite, infilled by barite, sphalerite and chalcopyrite. Dendrite growth textures branch outwards towards the chimney wall. The middle of the three zones comprises largely pyrite with chalcopyrite and sphalerite as thin veinlets and infillings. The innermost zone is dominated by chalcopyrite. Minor gold and bismuth tellurides occur at the boundary between the middle and the inner zones. The inner zone is interpreted as the hydrothermal conduit lining and in all cases is defined by bladed chalcopyrite, which shows a texture consistent with growth toward an inner open space. The central part of the conduit is now infilled with sphalerite or in one case pyrite, chalcopyrite, silica and minor barite. Fluid inclusion studies indicate the presence of high-temperature ( > 350°C) fluids of around 3.5% equivalent NaCI in the basal parts of the massive sulphide mound with cooler temperatures ( < 100°C) recorded in barite from the upper part of the mound. Barite is associated with chimney fragments, as worm tube infillings and as later cross-cutting veins. Preliminary S and Sr isotope data from sulphides and sulphates supports both igneous and seawater sources for sulphur in the hydrothermal system with evidence for seawater circulation and sulphate precipitation beneath the sulphide mound. The results are consistent with a similar model of chimney growth to that proposed for modern vent sites. It is proposed that a high temperature fluid flux through open conduits fed black smoker activity accompanied by lateral fluid diffusion through the chimney wall which mixed with seawater. The result of this is a combination of conductive cooling and fluid mixing leading to precipitation of distinctively zoned mineral assemblages across the vent conduit wall. The occurrence of tellurium, bismuth and precious metal-bearing phases indicates some similarities with the geochemistry of sulphides from other Palaeozoic massive sulphide deposits associated with felsic volcanic centres.

Sulphide chimney spires are striking features of m o d e r n seafloor massive sulphides. They are definitive evidence o f the venting of hightemperature hydrothermat fluid, which mixes with cold seawater and rapidly chills to precipitate sulphide and sulphate minerals. In the geological record, such features are scarce, with c o n v i n c i n g description limited to M e s o z o i c sulphides in Cyprus (Oudin & C o n s t a n t i n o u 1984), O m a n ( H a y m o n et al. 1989) and more recent K u r o k o deposits (Scott 1981). This is despite the suggestion t h a t 'black s m o k e r ' seafloor exhalative activity may date back to the Archaean (Vearncombe et al. 1995). However, it is p e r h a p s u n s u r p r i s i n g given the l o c a t i o n o f these features at the s u l p h i d e

m o u n d - s e a w a t e r interface where cessation o f h y d r o t h e r m a l activity leads to rapid erosion and oxidation of any sulphides formed. However, exceptionally preserved volcanogenic massive sulphide ores at Y a m a n Kasy, in the southern Urals have yielded fragments o f black smoker chimneys, clearly associated with a distinctive vent fauna (Little et al. 1997). A more detailed description o f the recovered vent fauna is contained in Little et al. (1998) and thus is not discussed here. The mineral associations, zonation and textures are very similar to black smoker chimneys recovered from active vent sites ( H a y m o n 1983; Tivey & M c D u f f 1990). Delicate textures defining the chimney wall and internal conduits are extremely well preserved

HERRINGTON, R. J., MASLENNIKOV,V. V., SPIRO,B., ZAYKOV,V. g. • LIttLE, C. T. S. 1998. Ancient vent chimney structures in the Silurian massive sulphides of the Urals. In." MILLS, R. A. & HARRISON,K, (eds) Modern Ocean Floor Processesand the GeologicalRecord, Geological Society, London, Special Publications, 148, 241-257

241

242

R.J. HERRINGTON E T AL.

and contrast in their zonation and morphology with sulphide-preserved biological tubes (Little et al. 1998).

,,loo okm ,

Miass

ht

map area

Geology and morphology of the sulphide deposit The Yaman-Kasy deposit is one of at least 60 volcanogenic massive sulphides discovered in the southern Urals. These are developed within a series of elongate, N-S trending structural zones, ranging in age from Ordovician to Carboniferous (Fig. 1). Yaman-Kasy is located in the Sakmara zone, a remnant of the Ordovician to Devonian-age marginal sea on the western side of the Uralian Ocean (Zaykov 1991). The Sakmara zone now forms an allochthonous unit, overthrust onto Permian rocks of the East European plate (Zonenshain et al. 1984; Savlieva & Nesbitt 1996). The massive sulphide deposit comprises an asymmetric, westward dipping lens of massive sulphides up to 37m thick and 90 to 100m in diameter (Fig. 2). The deposit contains c. 2.5 Mt of sulphide ore grading 2.5% Cu, 5.5% Zn, 3gt -1 Au and 33g t -1 Ag. The footwall of the orebody comprises highly altered rhyolites, which now take the form of quartz-chlorite-sericite altered volcanics that broadly underlie the deposit to the southwest. The immediate footwall to the massive sulphides comprises brecciated and veined altered rhyolite which is interpreted as the stockwork feeder to the massive sulphide lens (Fig. 3). This feeder zone sulphide comprises largely pyrite, chalcopyrite and silica. This passes up into a massive sulphide matrix supported by altered rhyolite breccia. The footwall also contains numerous thin (< 5 cm wide) later barite-pyrite :5 chalcopyrite veins. Brecciated texture, complex banding and crosscutting veining are the most common textures and there is little evidence for simple sulphide layering in the lens itself. Silica, the other significant phase is present largely as later cavity infillings. The sulphide lens shows a broad geochemical-mineralogical zoning upward from a pyrite rich lower part, through sphaleritechalcopyrite rich ore ( Z n : C u 3 : 1), to a chalcopyrite-sphalerite zone ( Z n : C u 1:1) (Zaykov 1991). The upper part of the lens is marked by increased brecciation and the presence of clastic layers of sulphides and oxidized sulphides, which are interpreted as the product of early seafloor alteration and reworking. Black, apparently organic-rich, sediment layers are found at the base of this horizon in some places. In the open pit, these layers of metal-rich clastic and

o

'osco KEY Massive sulphide deposits referred

]!! / ~-~'J

to in text Marginal allochthons Inter-arc sequences { Island arc sequences

52N

t

t

Sakrnar~ \ zone ',

-N =

\

I

Blyava '~' Yaman~

Kasy 100 km

",, ',

Palaeozoic arc and marginal sea complexes of the southern Urals Fig. 1. Geological map of the southern Urals showing broad structural units and settings of the massive sulphides. chemical sediments can be traced some distance away from the sulphide lens itself (Fig. 2). In turn, the mineralized sequence is overlain by a series of extrusive lavas and pyroclastics of rhyolitic composition, which are largely unaltered. The upper parts of the sulphide lens are mainly sulphide-cemented breccias formed of sphalerite-pyrite-barite and chalcopyrite-sphalerite-pyrite. These breccias are the main host to the vent fauna recovered from the deposit (Little et al. 1997), and to three of the recovered vent chimney fragments described in this paper. The fourth fragment was recovered from the wellbedded clastic sulphide material that forms the lateral equivalent of the massive sulphide lens at the southern end of the open pit. This unit shows distinctive grading of sulphide grains and is interpreted as having formed during the degra-

SILURIAN VENT CHIMNEYS

243

/

Plan

Rhyodacite to basalt lavas and hyaloclastites

..........

.... Rhyodacite lavas and hyaloclastites and intercalated siltstones

I-7~ --~ Sericite-quartz altered rhyo-dacite Line Sect

1L..,,,

if!i!

i!!il. Andesiticbasaltandbasalt

/ !iiiiiiiiiiill

Massive sulphide Projection of mined orebody Faults

Section 400 metres as.l. 3O0

Cherty ironstones Location of vent chimney debris I"

.-" i Outline of open pit

2O0 Fig. 2. Geological plan and section of the Yaman-Kasy orebody (after Zaykov 1991). dation, hydraulic transport and redeposition of material from the sulphide mound (Fig. 4). This would indicate that there was some original topographic relief to the sulphide mound, which is not unreasonable given the relief of such modern sulphide deposits like TAG (Humphris et al. 1995).

Sampling Fragments of vent zoned chimney conduit have been recovered from the Yaman Kasy deposit along with other irregular tubular structures which are likely to be partial fragments of chimney. Although in detail quite different in texture to one another, the broad features and zonation patterns of these fragments are quite similar. Three of these fragments are described here. Fragments 1 and 2 were recovered from the main sulphide mound (see Location 2 on Fig. 2) in material which contains abundant vent fossils (worm tubes, monoplacophorans and brachiopods). Fragment 3 was recovered from the

flanking detrital sulphide ore at the lateral margin of the sulphide mound (Location 3 on Fig. 2). Chimney fragment 1 was selected for sulphur isotopic analysis of both sulphides and barite. In addition, further samples of sulphides and barite were collected from the main sulphide body, initially for sulphur analysis. Subsequent to this, two of the barite samples were selected for ~7Sr/S6Sr determination (see Table 3). Barite and quartz are common phases throughout the sulphide body and are often suitable phases for fluid inclusion study. After preliminary microscopy, two samples of quartz and barite (YK1-05 and YK1-03) containing primary, negative crystal inclusions were prepared for fluid inclusion microthermometry.

Methodology Chimney fragments 1, 2 and 3 have been examined by normal reflected light microscopy and by SEM/microprobe. Samples were initially

244

R. J. HERRINGTON E T AL.

Fig. 3. Altered brecciated rhyolite footwall below massive sulphide body at Yaman-Kasy. Dark areas of massive pyrite-chalcopyrite infill brecciated and sericitized volcanics.

viewed in reflected light before preliminary qualitative analysis using an Hitachi $2500 SEM fitted with an EDS X-Ray analyser at The Natural History Museum (NHM). Fragment 1 was selected for more detailed microprobe study using analysis by a Cameca SX-50 microprobe fitted with a WDS X-Ray analyser which is also located at the NHM. Operating conditions for the Cameca SX-50 microprobe were 15kV and 15nA, calibrated with wellcharacterized standards at identical operating conditions immediately prior to analytical runs. The various calculated detection limits ( + 3o- of the background) are indicated in Table 1. Fluid inclusion heating and freezing measurements of

~ 8 0 # m thick doubly polished wafers were carried out using a Linkham THMS600 stage at Imperial College. This stage is calibrated regularly with synthetic inclusion standards and an applied correction equation is calculated at regular intervals (normally 3-monthly intervals since there is negligible instrument drift). Measurement precision is +0.1°C on homogenization and last melting measurements with an accuracy of +0.2°C in the range - 6 0 ° C to + 30°C and ±0.5°C above + 30°C. Hand picked and precision drilled sulphides and sulphates were prepared for sulphur and 87Sr/86Sr isotope determination at the N I G L - N E R C facility, Keyworth. Sulphides

SILURIAN VENT CHIMNEYS

245

Table 1. EPMA Analyses of pyrite along profile across chimney fragment 2, sample YKI-03 (see Fig. 13for location

of analyses). Analysed using Cameca SX-50 at The Natural History Museum (15kV, 15nA). BL = Below calculated limit of detection Analyses in weight % element Analysis point S

Fe

Cu

Zn

As

Cd

Sb

Te

Hg

Total

Calculated detection limit 1 2 3 4 5 6 7 8 9 10 11 12 13 14 16 17 18 19 20 21 22 23

0.09 46.72 46.92 46.52 46.69 46.69 46.67 46.86 45.73 45,90 46.71 46.53 46.70 44.99 45.85 45.49 46.79 47.18 47.04 47.53 47.45 46.86 46.59

0.10 BL 0.10 BL BL BL BL 0.23 0.28 0.18 BL BL 0.29 0.20 0.10 0.13 0.33 0.18 0.12 BL BL BL BL

0.18 BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL

0.06 BL BL 0, l I BL 0,06 BL BL 0.10 0.11 0.10 0.08 0.14 0.15 0.17 0.61 BL BL 0.13 0.13 0.06 0.13 BL

0.05 BL BL 0.07 BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL BL 0.07 BL

0.05 BL BL BL BL BL BL BL BL 0.10 BL 0.05 BL BL BL BL BL BE BL BL BL BL BL

0.09 BL BL BL BL BL BL BL 0.29 0.21 0.09 0.17 0.38 0.35 0.36 0.16 0.11 BL BL BL BL 0.12 0.30

0.05 0.07 0.13 BL BL BL BL 0.10 BL 0.10 BL 0.06 BL BL 0.07 0.05 BL 0.12 BL BL 0.08 0.13 0.11

99.99 100.50 100.51 99.94 99.87 99.57 100.42 96.28 97.51 99.88 99.18 98.00 96.38 97.28 98.88 100.14 100.41 100.52 100.61 100.82 100.13 100.16

0.08 52.97 53.21 53.67 53.16 53.00 52.69 52.98 49.77 50.86 52.86 52.12 50.45 50.62 50.7 52.31 52.76 52.81 53.14 52.87 53.11 52.75 53.06

Fig. 4. Layered clastic F e - C u sulphides (bright bands) and iron oxide-silica (darker bands) sediments, distal equivalent to the massive sulphide lens. Graduations shown in millimetres. Sulphide clasts show clear evidence for grading.

246

R. J. H E R R I N G T O N E T AL.

Fig. 5. Sulphide vent chimney spire fragment 1 in long section (slice shown in plan in Fig. 6b). Figures on scale are centimetres.

Fig. 6. Plan view of section through the chimney shown in Fig. 5. White ~V' marks show outer wall of spire, now overgrown by sulphides. Zones A, B, C, D in white are explained in the text. Box marked 'Scan area' is the area of Fig. 12 microprobe profile.

SILURIAN VENT CHIMNEYS

247

Table 2. Analyses of tellurium bearing phases from Yaman-Kasy ch#nney material, YKC3 and YKC4 (Chimney fragment 2 and 3, respectively). Analysed using Cameca SX-50 microprobe @ 15 kV 15 nA at The Natural Histot3, Museum Sample point no. phase

YKC3-01 39-1 Altaite

38-1 Altaite

38-2 Altaite

37-2 Altaite

37-3 Altaite

YKC4-03 13-1 Telluro-bismuthite

Pb Te Ag Fe Cu Bi S Total

57.80 35.73 < DL 2.77 2.17 < DL 0.66 99.13

58.10 36.04 < DL 2.84 2.14 < DL 0.76 99.88

57.50 35.70 < DL 2.23 1.53 < DL 0.34 97.36

59.73 37.59 < DL 2.09 2.39 < DL < DL 101.8

59.64 37.15 < DL 2.41 1.96 < DL < DL 101.6

! .3 ! 46.65 0.62 0.27 < DL 50.53 0.15 99.53

were analysed for sulphur following the methods of Robinson & Kuskabe (1975). Barite sulphur analyses followed the method of Coleman & Moore (1978). Isotopic determinations were carried out using a VG SIRA 10 mass spectrometer. Results are expressed as deviations from the Canyon Diabolo Troilite Standard expressed in permil (t534S%oCDT). The method has an overall analytical reproducibility of +0.1%o. 87Sr/S6Sr determinations were performed on a Finnigan M A T 262 mass spectrometer. The NIST NBS 987 standard gave an S7Sr/86Sr of 0.710201 (+0.000014 2or).

M i n e r a l o g y o f chimney material Fragment 1. This is the best studied sample, since a short length of spire measuring some 9 cm long with an oval cross-section with lateral dimensions 5 cm by 3.5 cm was recovered. A layer of fine sooty pyrite with layers of iron oxides defines what is believed to be the outer wall. One side of the spire is attached to further sulphide material, which is either a later overgrowth or else forms the matrix cementing the spire fragment (Figs 5 & 6). In cross-section the spire is clearly zoned in four broad mineralogical domains. Zone A, the outermost zone, is the most complex. It comprises dominantly pyrite, barite and chalcopyrite present as an anastomosing pyrite network infilled with dendritic and collomorphic barite and chalcopyrite. Chalcopyrite and barite also form cross-cutting microveinlets. The collomorphic barite occurs in the very outermost layer where it is seen to infill tubular openings between the pyrite network. SEM imaging of the outermost shell of the spire confirms that the barite fills roughly tubular

cavities between a honeycombed pyrite network. The innermost portion of this zone comprises largely pyrite which surrounds common barite and much rarer sphalerite 'dendrites' which have developed in places, with the branching always splaying towards the outside of the spire (Fig. 7). Zone B is dominated by fairly massive pyrite, which contains minor chalcopyrite, largely as infillings to cavities in the pyrite. Zone C is a distinctive but irregular massive chalcopyrite layer, which often shows a thin silica-rich boundary layer with zone B. This layer also contains many inclusions of sphalerite. In this layer are many tiny grains of telluride minerals, dominated by altaite (PbTe) and an unresolved complex A g - P b - T e phase. Very rare galena grains ( < 5 #m across) are also found in this zone. Zone D comprises a layer of pyrite, chalcopyrite and silica. Pyrite is present as granular pyritohedra and cubes in a matrix of chalcopyrite and silica. The pyrite grains range from 0.2 to 0.5 mm across. A distinctive thin silica band marks the boundary between Zone C and Zone D and it is interpreted that Zone D is a later infilling inside the hollow conduit lined by Zone C.

Fragment 2. This fragment is strikingly zoned in cross-section and measures approximately 5 cm in diameter (Fig. 8). The outermost region of the fragment is composed of rather vuggy and laminated pyrite with an extreme outer layer of partially oxidized pyrite, haematite and minor barite. A dominantly pyrite with minor chalcopyrite zone is next. Inside this zone is a chalcopyrite dominated zone which contains minor bismuth tellurides. Inside this, bladed pyrite crystals are developed, probably after marcasite. A massive chalcopyrite zone is next

248

R. J. HERRINGTON ET AL.

Fig. 7. Reflected light view of dark barite dendrites in pyrite-chalcopyrite matrix from Zone A on Fig. 6. ba = barite, py = pyrite, cpy = chalcopyrite.

Fig. 8. Zoned chimney fragment 2. Sp = sphalerite, cpy = chalcopyrite, py = pyrite. Graduated scale shows millimetres.

which passes inward into a zone of bladed chalcopyrite with interstitial sphalerite. Finally, the centre of the fragment is marked by massive collomorphic sphalerite, which shows extensive chalcopyrite disease. The inner 'chalcopyrite' zone actually contains two copper-bearing sulphide phases, which in places show complex lamellar intergrowths. In one grain this takes the form of an orthogonal tartan pattern of bright and dark zones when viewed by back-scattered electron microscopy (Fig. 9). The brighter phase is shown to be chalcopyrite whilst the darker phase appears to be a complex intergrowth of chalcopyrite and apparently isocubanite with an approximate formula Cu0.65 0.75Fe0.9S2. The tartan intergrowth clearly mimics the cubic symmetry of isocubanite.

structure, well defined by mineralogical zones, and measures approximately 10 cm by 4 cm (Fig. 10). The outer zone (zone 1 on Fig. 10) comprises layered collomorphic pyrite, often with tiny included grains of barite. Inside this is a chalcopyrite-rich zone (zone 2, Fig. 10) which is of variable thickness, up to 1 cm thick in places, which has a sharp inner contact with a mixed pyrite and chalcopyrite zone (zone 3 on Fig. 10). A thin, sooty pyrite layer that contains occasional telluride phases marks the sharp contact inside this chalcopyrite zone. Inside the mixed chalcopyrite-pyrite layer is a broader chalcopyrite-pyrite zone (zone 4 on Fig. 10) which contains distinctive bunches of tellurobismuthite lamellae (Fig. 11 and Table 2). The inner part of zone 4 is marked by bladed pyrite laths, which appear to mimic former marcasite which passes to granular pyritohedra infilled by interstitial chalcopyrite and silica. Finally in zone 4, laths of chalcopyrite occur, interstitially infilled with sphalerite. These laths all face into

Fragment 3. This fragment has the largest cross-

section of any of the chimneys yet found. In section it forms an embayed closed loop

249

SILURIAN VENT CHIMNEYS

Table 3. Sulphur and strontium isotope datafrom Yaman-Kasy Sample number

Description

Phase

(534S

YK1-01

Fossiliferous mound sulphide

matrix pyrite

+ 0.4

matrix sphalerite sphalerite infilling worm tube pyrite core chalcopyrite pyrite outer layer sphalerite pyrite barite pyrite

+ 0.2 + 0.5 + 1.4 + 2.0 + 2.9 + 0.3 + 1.6 + 0.8 + 0.0

YK 1-03

Chimney fragment 1

YK1-05

Massive barite-rich ore

95-B-1

Barite infilled pyritic worm tube

barite 95-B-3

Massive sulphide with vestimentiferan tubes

YKI 5

Clastic sulphide fragments with vent chimney

YK63 YK 107

Sphalerite-rich ore Clastic sulphide layers

YK1-04

Late footwall barite vein in massive pyrite alteration of volcanics

Detailed mineral chemistry study of chimney fragment A detailed microprobe study was m a d e of pyrite,

+ 0.6

chalcopyrite

+ 0.4

sphalerite pyrite pyrite pyrite layers on vestimentiferans

+1.0 +0.0 -0.7 -1.6 2.5 +0.5 + 25.8 +0.3 +0.2 -0.1 -0.3 -0.5 +0.9 +0.9 +1.2 + 26.7

barite infilling to tube pyrite

sphalerite pyrite chalcopyrite barite pyrite host pyrite at vein margin

the chimney cavity (zone 5 on Fig. 10) which is n o w infilled by sphalerite. This sphalerite occasionally shows o r n a m e n t a t i o n consistent with p s e u d o m o r p h after wurtzite. The sphalerite shows distinctive chalcopyrite disease texture and these two phases infill the former open conduit. This conduit appears as if it may have had a complex history since the presence of two annular chalcopyrite-rich zones suggests that at least two periods of chalcopyrite deposition occurred.

S7Sr/S6Sr = 0.70694

~7Sr/86Sr = 0.70933

+ 0.5 + 2.2

which is present t h r o u g h o u t the wall of the chimney spire fragment 1. This traverse is shown in Figs 6 and 12, chosen where the outer baritebearing zone (Zone A) and pyrite zone (Zone B) of the chimney are unusually thin, being directly developed over the chalcopyrite-rich core material (Zone C). A back-scattered electron image of the microprobe profiles is shown in Fig. 12 and the results are presented in Table 1. Pyrite in parts of the profile has analytical totals below 98% (both Fe and S totals are decreased). These low values are readily reproducible and are a function of the highly porous fine-grained pyrite present in this zone. Analysis was carried out as a closed run of points from 1 to 23, pyrite either side of the zone with low totals having totals close to 100%. The results for Cu and Te show distinctly elevated contents of up to 0.32 and

250

R.J. HERRINGTON E T A L .

Fig. 9. Tartan pattern of chalcopyrite (light grey) and isocubanite-chalcopyrite intergrowth in fragment 2. cpy = chalcopyrite.

0.38 wt%, highly enriched with respect to other zones where pyrite is relatively pure FeS2. Although the highest values of Cu and Te occur in grains where analytical totals are below 98%, significantly enriched Te and Cu is shown in adjacent areas (e.g. points 11 and 16) to support the previous statement. When the results for the outer wall of fragment 1 are plotted as profiles (Fig. 13), it is clear that both Cu and Te show values below detection limit close to the chimney margin. In this region, significant sphalerite and barite dendrites occur together with laminated layers of pyrite. Pyrite showing C u T e enrichment forms a band within the central zone (Zone B). The low Cu and Te values close to the inner zone appear to reflect the presence of discrete Te and Cu-bearing phases in the inner core zone which can accommodate these elements. Tellurides are absent from the outer zones of all of the chimney fragments studied to date and thus the decline in Te and Cu content in pyrite probably relates to the likely reduced mobility of these elements across the steep geochemical and temperature gradients across the vent chimney. The fact that this apparent zonation pattern of tellurium-bearing phases is seen in many of the sulphide chimney fragments, together with the

Fig. 10. Zoned chimney fragment 3. Zones marked 1 to 5 referred to in the text. Graduated scale shows millimetres.

251

SILURIAN VENT CHIMNEYS

Table 4. PrelhninaO, )quid inclusion analyses from Yaman Kasy chimney and sulphide mound material Sample number

Phase

T First melt

YK1-05

Quartz in massive mound pyrite

-28.1

YK 1-03

Quartz infilling chimney

T Last melt -1.9 -2. l -1.7 -3.1 2.3 -1.8 -1.9 -1.6 1.8 -1.2 1.4 -l.7 -1.6 -2.2 -1.4 -2.3 -1.9 -1.7 -2.5

362 371 330 345 340 337 335 335 327 341 338 310 315 315 321 320 304 228 205

-2.7 1.7 -1.8 -2.1 -1.3 -2.2 -2.0 -1.2 -2.1 -2.0 -1.4

280 279 276 261 241 232 181 187 190 162 160 155 126

1.2 -1.4 -1.6 2.1 -1.8 -1.4 -2.0 -1.3 -1.8 -2.0 1.3 -1.7 -1.5 2.O 1.6 -1.8 1.9

180 187 150 155 162 167 171 128 141 148 103 105 116 118 120 122 122

1.7 1.8

Barite Infilling chimney

evidence that the sulphide material a p p e a r s n o t to have been recrystallised (delicate relict textures in c o p p e r - b e a r i n g sulphide phases), w o u l d s u p p o r t the i n t e r p r e t a t i o n o f this z o n a t i o n as a p r i m a r y feature.

T Homogenization

Isotope chemistry T h e results o f the s u l p h u r isotopic analyses m a d e o n pyrite, chalcopyrite, sphalerite a n d barite f r o m vent c h i m n e y f r a g m e n t s a n d m o u n d

252

R.J. HERRINGTON ET AL.

Fig. 11. Back-scattered electron image of tellurobismuthite lamellae developed in porous pyrite zone in fragment 3.

material are shown in Table 3. Values of ~348 for sulphides lie in the range - 2 . 5 to + 2.9%0 whilst the limited barite data are divided between approximately + 26 and + 0.6%0. In detail, data from sulphides from chimney fragment 1 show a range of ~34S from +0.7 to +2.9 with the highest value from the outer portion of the chimney. Pyrite layers on fossil worm tubes show slightly negative (~348 (--2.5 to --1.6) compared to the chimney material. The two barite types indicated by the sulphur data were then analysed for 87Sr/S6Sr. The sample from the worm tube infill in the massive sulphides (95-B-1) gave a value of 0.70694 whilst the late cross-cutting footwall veining (YK1-04) gave a value of 0.70933.

Fluid &clusion study Fluid inclusion data from the two samples where inclusions could be measured are shown in Table 4 and on Fig. 14. The homogenization data from sample YK105, from the basal part of the sulphide mound,

confirm that high temperature fluids ( > 350°C) were trapped in the quartz. These temperatures are uncorrected for pressure. No evidence of boiling is recorded in the inclusions which means that if a hydrostatic head of seawater is assumed, a seawater depth of c. 1600m is indicated. Data from sample YK1-03 are measured in quartz and barite which infills the chimney conduit in fragment 1. These data record a lower fluid temperature, with some inclusions homogenizing as low as 100°C. Homogenization temperatures from quartz samples span the full temperature range observed, although the barite would appear to have formed at a lower temperature. Inclusions in barite are sometimes susceptible to leakage, however, loss of part of the inclusion fluid phase by leakage would serve to raise observed homogenization temperatures. This is not supported by the data, which indicate homogenization temperatures equal to or lower than quartz, which generally suffers no leakage problems. The last melt temperature data indicate only one broad population for fluid salinity. Using the equation developed by Bodnar (1993), ice

SILURIAN VENT CHIMNEYS

/

/

/

!

253

1 mm

Fig. 12. Back-scattered electron image of section of chimney wall shown in Fig. 7. Zones A, B, C referred to in the text. Analysis points shown are shown in Table 1 and on Fig. 14.

melt temperatures indicate a fluid salinity for all the samples in the range 2.1 to 5.1 wt% NaC1 equivalent, with a mean of 3.2 wt% NaC1 equivalent.

Discussion C o m p a r i s o n s with m o d e r n chimne?'s

Studies of m o d e r n vent chimney material (various) and actual in situ experiments at vent sites (Tivey 1995) provide very good models for vent chimney growth applicable to fossil examples. Haymon's (1983) original model proposes that initial high-temperature fluid venting would result in formation of an anhydrite spire. Temperature increase inside the hollow spire leads to chalcopyrite growth as a lining to the anhydrite. This very simple model is confirmed by later studies (Tivey & McDuff 1990). All the chimney fragments from Yaman-Kasy show the clear development of a chalcopyrite-rich layer that formed the inner lining of the formerly open conduit. Herzig et al. (1993) stress that the chimney wall may form a porous network with both conductive cooling and fluid mixing processes through the wall. This type of porous network is similar to the honeycomb pyrite layer seen on the outer part of fragment 1 from Yaman-Kasy. Isocubanite in modern chimneys is testament to the likely quenching of a high-temperature fluid in the vent. lsocubanite and chalcopyrite

intergrowth forms from the breakdown of an initial phase (iss) during rapid cooling. Quenching rapidly from temperatures of more than 210°C preserves the cubic form of isocubanite which would otherwise recrystallise to the stable phases chalcopyrite (Cabri et al. 1973). The occurrence of relict isocubanite texture in fragment 2 is unique in fossil vent material to our knowledge and is further testament to the unmetamorphosed nature of the Yaman-Kasy orebody (Little et al. 1997). In fragment 2, there is a distinct change in the mineralogy of the sulphides in the chimney wall coincident with the occurrence of bladed pyrite crystals, interpreted as having formed after marcasite. Marcasite is also seen in the outer laminated sulphide layers. Marcasite normally forms instead of pyrite at pH < 5 (Murowchick & Barnes 1986) and at temperatures below 200°C, indicating that fluid temperatures in this part of the chimney were significantly lower than the conditions inside the central conduit and were of significantly lower pH. Barite is present in this section of the chimney suggesting the fluids may also have been more oxidized. Dendrites of both sphalerite and barite have been recorded in the outer layers of modern chimney walls (Fouquet et al. 1993; Paradis et al. 1988). Similar dendrites are also developed in iron disulphides at modern vent sites (Duckworth et al. 1995) and these structures are taken as evidence of rapid cooling of hydrothermal vent fluid and high rates of nucleation and mineral growth (Fouquet et al. 1993). Paradis et

254

R.J. HERRINGTON ET AL. 0.4 ,~

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Fig. 13. Microprobe traverse across Fig. 13. Lines shown are 3 point moving averages for copper (solid) and tellurium (dashed) in weight % element.

Barite Quartz E 5 ®

E

E

13 Z

0 100 200 300 400 Homogenizationtemperaturein C Fig. 14. Fluid inclusion histogram of measurements

from quartz and barite from the mound and chimney sulphide material (measurements not corrected for pressure - see Table 4).

chalcopyrite dominated inner zone. In the case of chimney fragment 4, these are abundant as apparently later infillings of tellurobismuthite in a zone that parallels the chimney outer wall. Native bismuth is recorded as a significant p h a s e in chimneys from the Escanaba trough where the massive sulphides are associated with sediment dominated sequences (Koski et al. 1988). Microprobe analysis of pyrite across the chimney wall of fragment 1 shows a zonation of Te and Cu, which relates to the broad mineralogical zones of the chimney wall. Similar Te zonation is recorded in modern chimney material recovered from T A G during the BRAVEX cruise in 1994 (I. Butler 1997 pers. comm.). Tellurium and copper are likely to have been introduced by the hydrothermal fluid and zonation of both elements in pyrite away from the conduit supports this hypothesis.

Fluid chemistry al. (1988) note that the dendrites always point outwards toward the chimney wall which is the direction of dendrite growth. Fragment 1 from Yaman-Kasy has dendrites of both sphalerite and barite in outer zones of the spire and these dendrites point outward. All the chimney fragments studied contain tellurides, all found in similar spatial positions in the chimney wall, namely at the outer limit of a

Fluid inclusion analysis indicates the presence of a high temperature (>350°C) hydrothermal fluid with a salinity similar to seawater in the deeper parts of the sulphide mound. There is no evidence for fluid boiling which suggests that the sulphides formed in at least 1600 m of seawater. Inclusion data further indicates a lower temperature fluid phase trapped in the upper part of the orebody, suggesting that the hydrothermal

SILURIAN VENT CHIMNEYS fluids were cooling and/or mixing at the top of the massive sulphide body. This scenario is broadly similar to what we know from modern vent sites, where hydrothermal fluid cools and mixes with cold seawater. Sulphur isotope analyses of sulphides from modern vent sites have been reviewed in Duckworth et al. (1995) and sulphides from YamanKasy fall largely within this range of 0 to + 6%. These low (~348 values are consistent with an igneous sulphur source, and an associated barite of similar low (~348 indicates that it probably formed from oxidised sulphur of similar source to the sulphide material. The limited sampling across the actual chimney material itself (sample YK1-03) suggests a general increase m b34S from core to outside ( + 1.4 to + 2.9). This is similar to patterns from some modern chimneys which have been sampled (e.g. Stryrt et al. 1981; Woodruff & Shanks 1988; Hannington et al. 1986; Shanks & Seyfried 1987). Elevated ~348 values are likely to be due to the increased component of reduced seawater sulphate in sulphides formed in the outer layers of the chimney (Woodruff & Shanks 1988) and such a model is proposed for the Yaman-Kasy chimney studied. However, a number of the pyrite samples show slightly negative 634S signatures• Negative sulphide values are rare in studies of modern vents (Duckworth et al. 1995). The bulk of the negative samples from Yaman-Kasy are directly associated with fossils, with the most negative value of - 2 . 5 coming from pyrite layers developed on the outside of vestimentiferan worm tubes. These layers contain what is interpreted as bacterial cell ghosts (Little et al. 1997) and it is proposed that the negative sulphur values may reflect a component of bacterially reduced seawater sulphate in these sulphide layers• In addition to the barite sample with low 634S, two samples gave values close to +26, which falls within the range for Silurian seawater sulphate which was +25 to +27 (Claypool et al. 1980). Suphur and strontium isotopes can be used to assess fluid mixing in fluids of distinctly different signatures (Styrt et al. 1981; Mills et al. in press)• At Yaman Kasy, the footwall vein barite with the seawater (~348 signature has a ~TSr/~6Sr value of 0.70933 which is close to the published value of 0.70880 for Silurian seawater (Burke et al. 1982). In contrast, the 87Sr/86Sr analyses of the barite with low (~348 had a 87Sr/S6Sr value of 0.70694, which is incompatible with a direct Silurian seawater source for strontium. These results support a model of two sources for sulphur and strontium in the Yaman-Kasy hydrothermal system, one compatible with a •

•

255

seawater origin and the other more compatible with an igneous derivation. In addition, these results indicate that seawater sulphate, apparently not mixed with a fluid of igneous source, circulated below the massive sulphide lens into the altered footwall after deposition of massive sulphides. A similar pattern of deep seawater circulation is seen from the 87Sr/86Sr data from anhydrite formed deep in the modern TAG system (Mills et al. 1998).

Conclusions Chimney fragments recovered from the massive sulphides at Yaman-Kasy show a broad 3-fold mineralogical zonation, similar to material recovered from modern extinct chimneys. An outer zone, of largely pyrite and marcasite formed the exterior wall. Honeycombed pyrite with sulphide and sulphate developed where vent fluid cooled and mixed with sulphate-rich seawater• Dendrite growth textures are evidence of rapid mineral growth from super-saturated fluids. A middle zone of largely pyrite, chalcopyrite and sphalerite passes to an inner core of chalcopyrite which represented the inner lining of the hydrothermal conduit, identical to the lining of modern chimneys• This inner lining often shows a bladed texture, evidence of growth into an open void, and a relict isocubanite texture is evidence of the precipitation of copper-bearing sulphides from quenched hightemperature fluids. Preservation of this original isocubanite-chalcopyrite intergrowth texture is unique for ancient vent material and is testament to the high degree of preservation of the YamanKasy material• These results are consistent with a similar model of chimney growth to that proposed for modern vent sites. It is proposed that a high temperature fluid flux through open conduits fed black smoker activity accompanied by lateral fluid diffusion through the chimney wall which mixed with seawater• The result of this is a combination of conductive cooling and fluid mixing leading to precipitation of distinctively zoned mineral assemblages across the vent conduit wall. Fluid inclusion studies confirm the presence of a high-temperature (>350°C) hydrothermal fluid of around 3.2% equivalent NaC1 which was cooled and mixed with seawater in the upper part of the sulphide mound. Preliminary S isotope data from sulphides is generally close to zero or mildly positive, indicative of derivation of the bulk of the sulphur from an igneous source, probably leached and transported by the

256

R . J . H E R R I N G T O N ET AL,

high t e m p e r a t u r e fluid f r o m footwall volcanic sequences. A single barite analysis shows a similar igneous s u l p h u r source. O t h e r barite samples give values o f a r o u n d + 26, c o m p a t i b l e with a Silurian seawater sulphate source. 87Sr/86Sr d a t a f r o m the t w o barite types s u p p o r t s the m o d e l o f two s u l p h u r sources, a n d the seawater signature f r o m the footwall vein barite is evidence o f seawater circulation a n d sulphate p r e c i p i t a t i o n b e n e a t h the sulphide m o u n d . T h e o c c u r r e n c e o f high levels o f tellurium, b i s m u t h a n d precious metals indicates s o m e similarities with the g e o c h e m i s t r y o f sulphides f r o m o t h e r Palaeozoic massive sulphide deposits a s s o c i a t e d w i t h felsic v o l c a n i c c e n t r e s (e.g. Eastern A u s t r a l i a - H u s t o n et al. 1995; N e v e s C o r v o , P o r t u g a l - P i n t o et al. 1994). This suggests that the Y a m a n - K a s y d e p o s i t m a y have m o r e in c o m m o n with this arc or rift-related geological setting t h a n m i d - o c e a n ridge settings like T A G . John Spratt is thanked for his help with microprobe analyses. The support of NERC grant GR3/10903 and Royal Society Joint Research Project funding for this research is acknowledged. The paper was much improved by comments from C. J. Stanley, J. S. Spratt, from the three anonymous reviewers and the volume editor (RAM).

References BODNAR, R. J. 1993. Revised equation and table for determining the freezing point depression of H20NaC1 solutions. Geochimica el Cosmochimica Acta, 57, 683-684. BURKE, W. H., DENISON, R. E., HETHERINGTON,E. A., KOEPN1CK, R. B., NELSON, N. F. & OTTO, J. B. 1982. Variation of seawater 87Sr/S6Sr throughout Phanerozoic time. Geology, 10, 516 519. CABRI, L. J., HALL, S. R., SZYMANSKI,J. T. & STEWART, J. M. 1973. On the transformation of cubanite. Canadian Mineralogist, I2, 33 38. CLAYPOOL,G. E., HOLSER,W. T., KAPLAN,I. U., SAKAI, H. & ZAK, I. 1980. The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chemical Geology, 28, 199 260. COLEMAN, M. L. & MOORE, M. P. 1978. Direct reduction of sulphates to sulphur dioxide for isotopic analysis. Analytical Chemistry. 50, 15941595. DUCKWORTH, R. C., KNOTT, R., FALLICK, A. E., R1CKARD, D., MURTON, B. J. ~,~ VAN DOVER, C. 1995. Mineralogy and sulphur isotope geochemistry of the Broken Spur sulphides, 29°N, MidAtlantic Ridge. Geological SocieO'. London, Special Publications, 87, 175 190. FOUQUET, Y., VON STACKELBERG, U., CHARLOU, J-L., ERZINGER, J, HERZIG, P., MUHE, R. & WIEDICKE, M. 1993. Metallogenesis in back-arc environments: The Lau Basin example. Economic Geol-

ogy, 88, 21542181. HANNINGTON, M. D., SCOTT, S. D., JOHNSON, I. R. & HALL, G. E. M. 1986. Sulfur isotopic evidence for local, non-equilibrium mixing during formation of a seafloor polymetallic massive sulfide deposit: Axial seamount. GAC-MAC Program with abstracts', 11, 77. HAYMON, R. M. 1983. Growth history of hydrothermal 'black smoker' chimneys. Nature, 31, 695 698. , KOSKI, R. A. & ABRAMS, M. J. 1989. Hydrothermal discharge zones beneath massive sulfide deposits mapped in the Oman ophiolite. Geology, 17, 531-535. HERZIG, P. M., HANNINGTON,M. D., FOUQUET,Y., VON STACKELBERG, U. & PETERSEN, S. 1993. Gold-rich polymetallic sulphides from the Lau back-arc and implications for the geochemistry of gold in seafloor hydrothennal systems of the southwest Pacific. Economic Geology, 88, 2182-2209. HUMPHRIS, S. E., HERZIG, P. M., MILLER, D. J. et al. 1995. The internal structure of an active sea-floor massive sulphide deposit. Nature, 377, 713 716. HUSTON, D. L., SIE, S. H., SUTER, G. F., COOKE, D. R. & BOTh, R. A. 1995. Trace elements in sulfide minerals from eastern Australian volcanic-hosted massive sulfide deposits; Part 1, Proton microprobe analyses of pyrite, chalcopyrite and sphalerite. Economic Geology, 90, 1167 1196. KOSKI, R. A., SHANKS, W. C. III, BOHRSON, W. A. & OSCARSON, R. L. 1988. The composition of massive sulfide deposits from the sediment-covered floor of the Escanaba trough, Gorda Ridge: hnplications for depositional processes. Canadian Mhwralogist, 26, 655 674. LITTLE, C. T. S., HERRINGTON, R. J., MASLENNIKOV,V. V., MORRIS, N. J. & ZAYKOV,V. V. 1997. Silurian hydrothermal-vent community from the southern Urals, Russia. Nature, 385, 14(~148. ,, , & - - (1998) The fossil record of hydrothermal vent communities. This volume. MILLS, R. A., TEAGLE,D. A. H & TWEY, M. K. (1998) Fluid mixing and anhydrite precipitation within the TAG mound. In." HUMPNRJS el al. (eds), Proceedings of the ODP, Scient(fic Results, TX (College Station), 158, 119-128. MUROWCHICK, J. B. & BARNES, H. L. 1986. Marcasite precipitation from hydrothermal solutions. Geochimica et Cosmochimica Acta, 50, 2615 2629 OUDIN, E. & CONSTANTINOU,G. 1984. ~Black smoker' chimney fragments in Cyprus sulphide deposits. Nature, 308, 349-353. PARADIS, S., JONASSON,I. R., LE CHEMINANT,G. M. & WATKINSON, D. 1988. Two zinc-rich chimneys from the Plume Site, southern Juan de Fuca Ridge, Canadian Mineralogist, 26, 637-654. PINTO, A., BOWLES,J. F. W. & GASPER,O. C. 1994. The mineral chemistry and textures of wittichenite, miharaite, carrollite, mawsonite and In-Bi-Hg tennantite from Neves-Corvo (Portugal). Abstracts oJ'the General Meeting of lMA, 16, 329. ROBINSON, B. W. & KUSAKABE,M. 1975. Quantitative preparation of sulphur dioxide for 34S/32S analysis from sulfides by combustion with cu-

SILURIAN VENT CHIMNEYS prous oxide. Analytical Chemistry, 47, 1179-1181. SAVL~EVA,G. N. & NESB~Tr, R. W. 1996. A synthesis of the stratigraphic and tectonic setting of the Uralian ophiolites. Journal of The Geological Society of London, 153, 525 537. SCOTT, S. D. 1981. Small chimneys from Japanese Kuroko deposits, h~. GOLDIE, R. & BOTRTLL,T. J. (eds) Seminars on Seafloor Hydrothermal Systems. Geoscience Canada, 8, 103-104. SHANKS, W. C., I I I & SEYFRIED,W. E., Jr. 1987. Stable isotope studies of vent fluids and chimney minerals, southern Juan de Fuca Ridge: Sodium metasomatism and seawater sulfate reduction. Journal of Geophysical Research, 92, 11 38711 399. STYRT, M. M., BRACKMANN,A. J., HOLLAND,H. D., CLARK, B. C., PISUTHA-ARNOND,V., ELDRIDGE,C. S. & OnMOTO, H. 1981. The mineralogy and the isotopic composition of sulfur in hydrothermal sulfide/sulfate deposits on the East Pacific Rise, 21°N latitude. Earth and Planetary Science Letters, 82, 36-48. Tlvwv, M. K. 1995. The influence of hydrothermal fluid composition and advection rates on black smoker. Geochimica et Cosmochimica Acta, 59(10), 1933 1949.

257

& McDuvv, R. E. 1990. Mineral precipitation in the walls of black smoker chimneys; a quantitative model of transport and chemical reaction. Journal of Geophysical Research, B, Solid Earth and Planets. 95(8), 1261~12637. VEARNCOMBE, S, BARLEY, M. E., GROVES, D. I., MCNAUGHTON, N. J., MJKUCKI, E. J. & VERANCOMBE, J. R. 1995. 3..26 Ga black smoker-type mineralization in the Strelley Belt, Pilbara Craton, Western Australia. Journal of the Geological Society of London, 152, 587-590. WOODRUFF, L. G. & SHANKS, W, C., III. 1988. Sulfur isotope study of chimney minerals and vent fluids from 21°N, East Pacific Rise: Hydrothermal sulfur sources and disequilibrium sulfate reduction. Journal of Geophysical Research, 93, 45624572. ZAYKOV,V. V. 1991, VolcanieiO' and Massive Sulphide -

-

Hills of Palaeooceanic Margins (Urals"and Siberia Massive Sulphide-Bearing Zones' .for Example). Moscow, Nauka publishing House. ZONENSHAIN, L. P., KORINEVSKI,V. G., KAZMIN,V. G., PECHERSKI, D. M., KHAIN, V. V, & MATVEENKOV, V. V. 1984. Plate tectonic model of the South Urals development. Tectonophysics, 109, 95 135.

The fossil record of hydrothermal vent communities C. T. S. L I T T L E 1, R. J. H E R R I N G T O N

1, V. V. M A S L E N N I K O V 2 & V. V. Z A Y K O V 2

1The Natural History Museum, Cromwell Road, London S W 7 5BD, U K 2 Institute o f Mineralogy, Urals Branch Russian A c a d e m y o f Sciences, Miass, Chelyabinsk district 456301, Russia

Abstract: There are 19 known fossiliferous volcanogenic massive sulphide (VMS) deposits which range in age from the Silurian to the Eocene. Most of these are in the Ural Mountains, Russia. The deposits contain assemblages of inarticulate and rhynchonellid brachiopods; gastropod, bivalve and monoplacophoran molluscs; and a small diversity of worm tube morphologies, some of which may be attributable to alvinellid polychaetes and vestimentiferans. The fossils are preserved mainly as external moulds of pyrite, which is consistent with biomineralization processes occurring at modern vent sites. Most of the fossil taxa are new to science, but the lack of original shells and organic tubes makes placement in existing phylogenetic schemes difficult. A comparison between modern vent communities and fossil vent assemblages shows that vestimentiferans, alvinellid polychaetes, bivalves, gastropods, monoplacophorans and perhaps brachiopods are shared at higher taxonomic levels, but that arthropods are found only in the modern communities. There are no direct ancestor-descendant relationships between the fossil and modern vent molluscs and brachiopods. This demonstrates that the modern vent environment is not a refuge for the known Palaeozoic and Mesozoic shelly vent taxa. Hence, taxonomic groups have moved in and out of the vent ecosystem through time. These findings are discussed in relation to alternative hypotheses for the origins of modern vent communities. That hydrothermal vent communities have a fossil record at all is fortuitous, considering their sporadic distribution on oceanic spreading ridges, the vagaries of fossilization at vent sites (discussed later), and how little Phanerozoic ocean floor has been obducted, without major metamorphism, and is now accessible for study. In this paper we review the fossil record of vent communities using published work and new data, and discuss the data with reference to competing hypotheses for the origin of the modern vent biota. Newman (1985) proposed that the modern vent environment is an extinction-resistant refuge and that m o d e r n vent communities contain a proportion of Mesozoic or even Palaeozoic 'relic' taxa. Based on this he suggested that vent communities offer 'a glimpse of antiquity' (see McArthur & Tunnicliffe this volume, for a fuller discussion). This intuitively attractive hypothesis makes two predictions: that modern and ancient vent communities should have similar taxonomic structures, and that there should be direct ancestor-descendant relationships between some of these taxa (Tunnicliffe 1991, 1992). An alternative hypothesis is that the vent environment has the potential for rapid bursts of adaptive radiation and frequent extinction events (Hickman 1984), reflecting its inherent

instability, both spatial and temporal (Van Dover 1995; Tunnicliffe et al. 1996; Tunnicliffe & Fowler 1996; Juniper & Tunnicliffe 1997). This predicts that modern and ancient vent communities should have fewer shared groups and fewer direct ancestor-descendant relationships than would be produced by Newman's proposal.

Fossil vent assemblages Macrofossils have been recovered from 19 VMS deposits to date (Fig. 1): one from the Eocene (Boirat & Fouquet 1986), one from the Palaeocene or late Cretaceous (Oudin et al. 1985), seven from the Cretaceous (Oudin & Constantinou 1984; Haymon et al. 1984: Haymon & Koski 1985), one from the Jurassic, six from the Devonian (Maslennikov 1991; Kuznetsov et al. 1988, 1991a,b; Zaykov et al. 1996), and three from the Silurian (Ivanov 1959; Kuznetsov et al. 1993; Little et al. 1997). Not considered here are a number of Palaeozoic fossil assemblages from sediment hosted (SEDEX) deposits (Banks 1985; Moore et al. 1986; von Bitter et al. 1992; Russell 1996; reviews in Campbell & Bottjer 1995 and Walter 1996) because their environment of formation was different to the mid-ocean ridge and arc settings of VMS deposits.

LITTLE, C. T. S., HERRINGTON,R. J., MASLENNIKOV,V. V. & ZAYKOV,V. V. 1998. The fossil record of hydrothermal vent communities. In. MILLS,R. A. & HARRISON, K. (eds) Modern Ocean Floor Processes and the Geological Record, Geological Society, London, Special Publications, 148, 259-270

259

260

C.T.S. LITTLE ET AL.

Safyanovka, Oktyabrsk, ~ i l e i n o e , BuriL)ay, " f ~ a , Sibay (L-I~ D; U) 'L Kapedhesi'~"~V

'~/'ff ~ 0Barlo (E; P)

2. "~ Figueroak (L J; P) ~ . ,

-

,~ r~-.~ "L....~-.~.~.

Sha~ ~ ,2(U K/P: P (K;T) ~ - ~ ( ~ ' ) Ocean system: I y. ° U = Uralian, P = Pacific, T = Tethys J " °

[ / ~(~

-,t.b

I

Fig. 1. Present geographical distribution of fossiliferous VMS deposits. Ages of the deposits and the ocean system in which they formed in brackets. P = Palaeocene, E = Eocene, K = Cretaceous, J Jurassic, D = Devonian, S - Silurian; L - Lower, M = Middle, U - Upper. =

Cenozoic and M e s o z o i c

The Barlo (Zambales ophiolite, Philippines; upper Eocene) and Azema (New Caledonia; Senonian or Palaeocene) VMS deposits contain fossil tubes (Table l) which are circular to ovoid in cross-section (Oudin et al. 1985: Boirat & Fouquet 1986). Oudin & Constatinou (1984) figured circular pyritic tubes (Table 1) in polished section from the Peristerka VMS deposit (Troodos ophiolite, Cyprus; upper Cretaceous, 91 Ma), but unfortunately did not give a scale for these. They compared the Barlo, Azema and Peristerka fossil tubes to silicified and pyrite overgrown tubes of the polychaete Ah, inella pompejana and other worms, not identified, from modern vent deposits on the East Pacific Rise (EPR) and Juan de Fuca Ridge (JdFR). Sinuous and coiled pyrite tubes, some with well developed flanges, have recently been found in pyrite blocks from the dumps of the Sha, Kapedhes, Memi, Kambia and Kinousa VMS deposits (all from the Troodos ophiolite, Cyprus; upper Cretaceous, 91 Ma.). In addition the last three deposits also contain specimens of a small, undescribed cerithioid gastropod while Memi also has specimens of another undescribed ?epitoniid gastropod (C. Little & M. Morisseau, in prep.). All the Cyprus vent fossils are hosted by black, fine grained pyrite and are associated with collomorphic layers. There is a variety of fossil tube morphologies

(Table 1) in the Bayda VMS deposit (Samail ophiolite, Oman; early upper Cretaceous, ~95 Ma) which formed on the crest of a spreading ridge (Haymon et al. 1989), perhaps in a marginal basin (MacLeod & Rothery 1992). The tubes are sinuous, closely spaced, but randomly orientated within a matrix of sphalerite and pyrite and are circular to oval in cross section (Haymon et al. 1984; Haymon & Koski 1985). Partial recrystallization in places destroys the fine details of the tube walls. Haymon & Koski (1985) found many similarities in preservation and mineral zonation between the Bayda tubes and tubes of Alvinella pompejana and the vestimentiferan ( = o b t u r a t e pogonophoran, Southward 1991) Riftia pach)Ttila entombed in sulphides at modern EPR 21°N and J d F R vents; however, they cautioned against direct taxonomic comparison with these species. The newly discovered Figueroa deposit (California, USA; lower Jurassic; C. Little & R. Haymon, in prep.) is hosted by pillow basalts incorporated within an accretionary prism. It contains a fossil assemblage consisting of the brachiopod Anarhynchia cf. gabbi and an undescribed gastropod species; pyrite and silica tubes with well developed flanges and wavy longitudinal ornamentation, morphologically similar to Ridgeia and Tevnia; and filamentous structures preserved in silica, similar in size and morphology to silicified microbial filaments reported by Juniper & Tebo (1995).

20 30 0.6 10 40 (72) 1

0.8-4.0

U. Eocene

L. Palaeocene/U. Cretaceous U. Cretaceous

U. Cretaceous

M. Devonian

M. Devonian

M. Devonian

M. Devonian L. Devonian L. Devonian L. Devonian

Silurian Silurian Silurian

Barlo

Azema

Peristerka

Bayda

Sibay large tubes

Sibay small tubes

Safyanovka

Uzelga Oktyabrsk Buribay Yubileinoe

Ljeviha Krasnogvardeyski Yaman-Kasy large tubes

Yaman-Kasy small Silurian tubes

75

300 1.5 400

30 4 50 50

47

70

400

17

Framboidal py

py py Arsenian py

Interlaminated py & qz py py Collomorphic py py

py

Interlaminated py (_<0.5 thick) & either qz or void py laminae (_<0.02 thick)+qz

PY

Collomorphic & dendritic py

Framboidal py

0.2 1.0

0.014).08

-

0.16 0.64

0.2-0.4


0.15

Wall thickness

py+sp+qz

ba q±qz or void

Smooth

Smooth

PY Sulphides (py ma s p ) - + b a i q z p y + s p ~ b a or void

Flanges Wavy longitudinal striations Internal cross-cutting annulations

cpy qz py

py+sp+qz+Fe carbonates

Flanges

-

Void or s p + q z

Ivanov 1959 Kuznetsov et al. 1993; Little et al. 1997 Kuznetsov et al. 1993; Little el al. 1997

Maslennikov 1991, this volume V.A. Prokin pers. comm.

Oudin & Constantinou 1984 Haymon et al. 1984; H a y m o n & Koski 1985 Kuznetsov et al. 1988, 1991a,b; Maslennikov 1991, this volume Kuznetsov et al. 1988, 1991a,b; Maslennikov 1991, this volume

Boirat & Fouquet 1986 Oudin et al. 1985

py ba + qz

Reference

Infilling mineralogy

Longitudinal ridges and concentric annulations

Some annulated

Ornamentation

All measurements in mm. Abbreviations: py = pyrite, cpy = chalcopyrite, ma = marcasite, sp = sphalerite, ba = barite, qz = quartz; U. = Upper, M. = Middle, L. = Lower. Dash (-) indicates no information available. 1Value in brackets refers to an anomalous large conical specimen.

5 1.0-1.5 6 7 5-6

1.0-7.5

0.4-4.0

2 10

1 5

1 2

~5

Age

Deposit

Diameter Maximum Wall mineralogy observed length

Table 1. Tube characteristics f r o m 13 fossiliJi, rous V M S deposits

262

C.T.S. LITTLE E T AL.

Fig. 2. Sulphide block from the Safyanovka deposit packed with pyritic tubes orientated roughly parallel to each other. Scale bar - 10ram.

Devonian

The Devonian of the Ural Mountains, Russia contains at least six fossiliferous VMS deposits (Fig. 1). Several small pyrite tubes (Table 1) were found in drillcores from the Yubileinoe (southern Urals; lower Devonian; ore bodies number 1 and 2) and Uzelga (southern Urals; middle Devonian: ore body number 4) deposits. Both deposits are hosted by felsic volcanics and Yubileinoe is thought to have formed in an intra-arc rift setting. Two small pyrite tubes (Table 1) have been found in the ore dump of the Buribuy deposit (southern Urals: lower Devonian) which is hosted by basalts and dacites and may have formed in an intra or inter-arc rift. The Oktyarbrsk deposit (southern Urals; lower Devonian) also has small pyrite tubes as well as poorly preserved moulds of indeterminate brachiopods or bivalves which are sometimes infilled by quartz (Maslennikov 1991). The deposit is hosted by felsic volcanics and formed in an intra-arc rift. Unfortunately, none of the fossils listed above is presently available for study, having been lost, mined away or made inaccessible in core stores. The Safyanovka VMS deposit (middle Urals; lower Devonian) is hosted by felsic volcanics which formed in an island arc setting, and now lies in the north part of the East-Magnitogorsk zone (Yazeva et al. 1992). The deposit contains parallel-sided fossil tubes (Fig. 2; Table 1), circular in cross-section, which occur as clusters of specimens orientated parallel to each other (C. Little et al., in prep.). The Sibay VMS deposit (southern Urals; middle Devonian) formed in a small rift (1-

3 km wide) cutting a large felsic volcanic dome complex located in an inter-arc basin, part of the Magnitogorsk West Mugodjarian zone (Maslennikov 1991; Zaykov et al. 1996). The deposit consists of four stacked massive sulphide lenses (Zaykov et al. 1996). The basal three of these contain in situ fossil assemblages of large tubes with flanges and smaller, smooth tubes (Ivanov 1947, 1959; Ivanov et al. 1960, ShcheglovaBorodina 1956; Kuznetsov et al. 1988, 1991a,b; Maslennikov 1991). A species of bivalve (Fig. 3a & b) co-occurs with the tubes in the third lens (Ivanov et al. 1960; Maslennikov 1991). The fossils are located in pockets of pyrite (often collomorphic), chalcopyrite and carbonates that have escaped recrystallization. Most of the ten known specimens of bivalve were found articulated in presumed life position. All the specimens are external moulds formed by a thin layer of pyrite that preserves impressions of growth lines and a wrinkled periostracum. The species was originally identified as M o d i o m o r p h a cf. mytiloides (Ivanov et al. 1960) but has subsequently been considered to be vesicomyid-like and similar to the modern genus C a l y p t o g e n a by Kuznetsov et al. (1988, 1991a, b). On-going taxonomic work (C. Little, in prep.) confirms that the species belongs to the family Modiomorphidae and is not a vesicomyid. The large Sibay tubes (Fig. 3c & d; Table 1) occur as clumps of individuals orientated approximately parallel to each other. They are presumed to be in life position because they cut sulphide depositionary layers at right angles (Kuznetsov et al. 1991a, fig. 1A). They are parallel-sided, mostly straight, but some specimens are slightly curved; all are circular in cross-

THE FOSSIL RECORD OF HYDROTHERMAL VENT COMMUNITIES

263

I I 1 I

! !

1

t

_l Fig. 3. Representative fossils from the Sibay deposit. (a) Articulated modiomorphid bivalve, side view of left valve. (b) Same specimen, dorsal view. Arrows indicate impressions of wrinkled periostracum. Scale bar = 10 mm. (e) Longitudinal polished section of large tube which is infilled by pyrite (black) and carbonate (white). (d) Camera lucida drawing of the same tube showing the laminated pyrite tube wall with flanges. Scale bar = 5 mm.

section. The walls bear saw-toothed or sometimes hook-shaped flanges, up 1.6 mm long, and between 0.5 and 5ram apart (Fig. 3d). These tubes have been interpreted by Kuznetsov et al. (1988, 1991a,b) to be fossil vestimentiferans, similar in morphology to the modern basibranch genera Ridgeia, Tevnia and Oasisia. The smaller tubes from Sibay (Table 1) are circular in crosssection and are found in clusters with random orientation (Kuznetsov et al. 1991a, fig. 1C) They have been interpreted by Kuznetsov et al. (1991a) to be polychaetes related to the modern vent family Alvinellidae. Sihtrian

The three Silurian fossiliferous VMS deposits are also located in the Urals (Fig. 1). Ivanov (1959) illustrated a flanged pyrite tube (Table 1) from the Krasnogvardeyski deposit (middle Urals) that is similar to the larger tubes from the Sibay deposit. This deposit is hosted by andesites and dacites and may have formed in an island arc setting. The Ljeviha deposit (middle

Urals) is hosted by mafic volcanics (dacites and basalts) and its south flank has yielded small pyrite tubes (Table 1). None of the material from Krasnogvardeyski or Ljeviha is presently available for study. The Yaman-Kasy deposit (southern Urals) is hosted by felsic volcanics and formed in a marginal back-arc basin which was part of the Sakmara zone allochthonous zone (Zonenshain et al. 1984; Savelieva & Nesbitt 1996; Herrington et al. this volume). The deposit is a lens of massive sulphide (Herrington et al. this volume) containing, in its upper part, an in situ fossil assemblage of an undescribed species each of lingulate brachiopod, monoplacophoran and gastropod molluscs, and two tube morphologies: large with longitudinal striations, and small and annulated (Figs 4 & 5; Table 1; Kuznetsov et al. 1993; Little et al. 1997). The fossils are hosted within a matrix of sulphide cemented breccias, largely formed of pyrite, chalcopyrite, and sphalerite, which also contain zoned veins, collomorphic pyrite, and vent chimney debris (Herrington et al. this volume). Fragments of the

264

C. T. S. LITTLE E T AL. |

Fig. 4. Representative brachiopod and mollusc fossils from the Yaman-Kasy deposit. (a) & (b) Lingulate brachiopods, both epoxy resin external moulds of pedicle valves. (e) (f) Monoplacophorans; (c) slightly oblique view of complete specimen; (d) apex of incomplete specimen, side view; (e) anterior view of incomplete specimen; (f) detail of longitudinal polished section, white arrows point to the laminated pyrite tube wall with large flanges. Scale bars (a) (f) = 10 ram. (g) & (h) Gastropod, apertural and apical views; scale bar = 0.5ram.

large tubes and brachiopods are also found in layered epiclastic sulphides flanking the deposit (Herrington et al. this volume). The brachiopods (Fig. 4a & b) and monoplacophorans (Fig. 4 ~ f ) often co-occur and are found loosely scattered or in groups of up to ten specimens; they are also closely associated with the fossil tubes. At least 126 reasonably complete m o n o p l a c o p h o r a n specimens and 80 mostly disarticulated brachiopods have been collected. Two gastropod specimens (Fig. 4g & h) have recently been discovered associated with a monoplacophoran. The brachiopods, monoplacophorans and gastropods are preserved as external moulds of pyrite; all original shell material is absent (Fig. 4f). (The brachiopods

and monoplacophorans were figured by Kuznetsov et al. (1993) as Calyptogena-like vesicomyid bivalves and possible archaeogastropods, respectively.) The large and small Yaman-Kasy tubes are c o m m o n in the deposit (i>250 collected specimens of each) and occur as single specimens or discrete clumps of approximately parallel oriented individuals. The small tubes (Fig. 5a: Table 1) are parallel-sided, circular in cross-section and have zoned mineral infillings. The large tubes (Figs 5c~e; Table 1) are nontapering and most show varying degrees of compaction prior to fossilization, being bent (Fig. 5c), twisted and having elliptical to almost flat cross-sections. However, a few of the tubes are straight-sided with perfectly circular cross-

THE FOSSIL RECORD OF HYDROTHERMAL VENT COMMUNITIES

265

/

a

~:

l

/ Fig. 5. Representative fossil tubes from the Yaman-Kasy deposit. (a) Small annulated tube (presumed polychaete in Little et al. 1997); scale bar---3 ram. (b) Large conical tube; scale bar = 20 ram. (e)-(e) Large tubes (vestimentiferans in Little et al. 1997). The specimen (c) is bent and the ornament of longitudinal striations and concentric wrinkles (white arrow) shown in higher magnification (d) follows around the folds in the tube indicating that the original organic tubes were both flexible and epifaunal. Scale bars for (c) & (e) = 10mm; (d) = 2 mm.

sections (Fig. 5e). The walls of the large tubes have an external ornament of wavy longitudinal striations (Fig. 5d), exterior and sometimes interior to which arc often developed microlaminated pyrite layers, 0.2-2.7 mm thick (Little et al. 1997, figs 2b d). The thickness of these layers appears to relate to the mineralogy of the matrix surrounding the tubes: the layers are thinner where this is sphalerite and thicker where it is pyrite and/or chalcopyrite. Adjacent to the tube walls the pyrite layers are penetrated by circular holes approximately 1 #m in diameter (Little et al. 1997, figs 2c&d). These were interpreted as the 'ghosts' of former microbial cells by Little et al. (1997). Two conical fossils have been found in the Yaman-Kasy deposit; the larger of these tapers from 72 mm diameter at the top to 13 mm over a length of 160 mm (Fig. 5b). These are presumed to be the most posterior parts of the large tubes. Little et al. (1997) identified the larger tubes as vestimentiferans, following Kuznetsov et al. (1993), based on similarities of size, gross morphology (flexible, non-tapering tubes with closed conical posteriors) and life habit (living epifaunally in vent settings) with the recent axonobranch vestimentiferan Riftia p a c h ) ~ t i l a

(described in Jones 1985). The closest modern analogue to the small tubes may be the annulated sulphide layers that have been observed on the inside of a few Ah, inella p o m p e j a n a tubes from the E P R (Haymon 1982). The small tubes were figured by Kuznetsov et al. (1993) as alvinellid-like polychaetes. Discussion Fossilization at vent sites

None of the fossil vent brachiopods and molluscs examined by us has any original shell material. All the specimens are external moulds formed by thin layers of pyrite which show details of periostracum, growth lines and ornament. Some of the pyrite layers forming the monoplacophorans from Yaman-Kasy (Fig. 4f) are particularly thick (>0.5 mm), becoming collomorphic and obscuring any ornament. The pyrite layers are often fractured and traversed by later veins (particularly at Yaman-Kasy where the veins are chalcopyrite and barite) indicating that they formed early in the diagenetic history of the vent deposits. There is no original organic material left in the fossil vent tubes; most are

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C. T. S. LITTLE ET AL.

formed by layers of interlaminated pyrite and silica or laminated pyrite, which preserve details of the tube wall ornament such as flanges and striations. These details are lost when the layers become thicker and collomorphic (e.g. Little et al. 1997, fig. 2b). Most of the tubes are external moulds, but it is possible that direct replacement of the original tube has also occurred. The small tubes from Yaman-Kasy and those from Barlo are formed by replacement of the original tube wall by framboidal pyrite. Processes occurring at modern vent sites give insights into how the fossil preservation style outlined above may have occurred. Rapid carbonate shell dissolution is characteristic of vents (experimental mean dissolution rates of 1401aria 1 for aragonite shells and 40ltma i for calcite shells have been recorded at Guaymas Basin by Lutz el ell. 1994) because of highly acidic or alkaline hydrothermal fluids, and the fact that many sites are below the carbonate compensation depth (CCD). Dissolution pits are common on living vent mollusc shells, particularly the older parts (gastropod apices and bivalve umbones) where the periostracum has been lost (McLean & Haszprunar 1987; Bouchet & War6n 1991; Hunt 1992; Allen 1993; Lutz el al. 1994). Therefore, mollusc shells probably have a short post-mortem residence time. However, the mineral coatings (mostly Fe oxides) that are seen on many vent mollusc shells (e.g. War6n & Bouchet 1986: Tunnicliffe & Fontaine 1987; Okutani 1990) may delay dissolution and give a greater chance of subsequent fossilization in the form of external moulds. The tubes of vent dwelling vestimentiferans may become overgrown by mineral crusts during life. Some Ridgeia piscesae tubes from JdFR, for example, have silica (C. Little, personal observation) and Fe-rich crusts, the latter probably mediated in part by bacteria (Jannasch & Wirsen 1981; Tunnicliffe & Fontaine 1987). Mineralization is also associated with alvinellid polychaetes (Gaill & Hunt 1991; Juniper et ell. 1986; 1992; Juniper & Sarrazin 1995). Direct replacement of organic layers by pyrite, and interlayer collomorphic pyrite and marcasite growth, has been seen in an Ah,inella pompeiana tube from the EPR (C. Little, personal observation). Parah'inella Sldh'ncola forms temporary muco-particulate tubes on the chimneys in the north-east Pacific, altering the substrate mineralogy from anhydrite to marcasite (Juniper et al. 1992; Tunnicliffe et al. 1993; Juniper & Sarrazin 1995). Alvinellid and vestimentiferan tubes (Alvinella pompejana, Riftia pachyptila and Ridgeia piscesae, where identified) are commonly found incorporated into the sulphides of grow-

ing vent chimneys and mounds (e.g. Haymon et al. 1984; Cook & Stakes 1995). The tubes are initially replaced and/or overgrown by concentric layers of barite and silica (Cook & Stakes 1995) or interlaminated silica and pyrite (Haymon & Kastner 1981; Haymon et al. 1984; Haymon & Koski 1985). These layers then act as templates for later, higher temperature sulphide mineral growth (including sphalerite and collomorphic marcasite and pyrite), and in places are completely replaced by chalcopyrite and wurzite. After the death of the animals and during the mineralization process, the tubes remain open, allowing the passage of hydrothermal fluids and giving substantial porosity to the enclosing sulphides, but they gradually become blocked by concentric layers of first collomorphic pyrite, then aggregates of Fe and Zn sulphides (Haymonet al. 1984; Haymon & Koski 1985; Cook & Stakes 1995). Based on the above observations we propose a fossilization process at the ancient vent sites that begins with the growth of pyrite (or an iron sulphide precursor) and silica layers on the outside of epifaunal organic tubes and carbonate shells, mediated wholly or in part by microbial activity. This forms a hard crust during the life of the organisms and/or shortly post-mortem. Subsequent carbonate shell dissolution and scavenging of the organics removes the original shells and tubes and the fossils are entombed within the silica/sulphate/sulphide matrix of the vent deposits. The tubes remain open for a period and act as conduits for waning hydrothermal temperature flux, becoming infilled with zoned minerals. Later diagenetic processes, such as veining and recrystallization, cause partial or complete destruction of the fossils in the base and middle of the growing deposits. Fragments of the matrix containing fossils are episodically reworked into flanking layered epiclastic sulphides (representing quiescent periods?). The fossilization process ends with the vent deposit being covered by lavas and/or sediments to prevent destruction by supergene processes. Vent j o s s i l identification p r o b l e m s

All the fossil vent brachiopods and molluscs are undescribed (C. Little, in prep.). Most appear to be new species. This mirrors the high level of taxonomic novelty at modern vents (Tunnicliffe 1991, 1992). The lack of shell material in the fossils hinders their placement in existing phylogenetic schemes because many of the taxonomic characters that are applied to fossil brachiopods and molluscs (e.g. hinge structures, muscle scars, shell composition) cannot be seen

THE FOSSIL RECORD OF HYDROTHERMAL VENT COMMUNITIES

267

lates and rhynchonellids), gastropods, bivalves, monoplacophorans and a small diversity of tube morphologies, some of which may be attributable to vestimentiferans and alvinellid polychaetes. The assemblages range in age from the Microbial filaments 1 Silurian to the upper Eocene (at least 370 Ma) Vestimentiferans and represent vent communities from three Polychaetes major ocean systems (Pacific, Tethys and Urals palaeo-ocean). Given the small number of data Arthropods points it is difficult to assess whether ancient Bivalves vent communities were as spatially heterogeneous in taxonomic structure as modern comGastropods munities, which vary on scales of individual vent Monoplacophorans fields to major ridge systems (Tunnicliffe & Fowler 1996; Tunnicliffe et al. 1996; Juniper & Brachiopods Tunnicliffe 1997). If ancient communities were Fig. 6. Taxonomic comparison between grouped as geographically diverse, the known fossil vent modern vent communities (data from Tunnicliffe assemblages may not be characteristic of other, 1991, 1992) and fossil vent assemblages (Cenozoic = coeval, as yet undiscovered, vent assemblages. Barlo, Azema; Mesozoic = Peristerka, Memi, Sha, This can only be determined following the Kapedhes, Kambia, Kinousa, Bayda, Figueroa; Devonian = Sibay, Safyanovka, Oktyabrsk, Buribay, discovery of more fossiliferous vent deposits. Uzelga, Yubileinoe; Silurian = Ljeviha, Krasnogvar- Additional faunal assemblages are almost cerdeyski, Yaman-Kasy). Black squares indicate presence. tainly present amongst economic and subQuestion marks and hatched pattern indicate identifi- economic VMS deposits which have little cation and ecological uncertainties, respectively. Note: metamorphic overprint. 1 includes the microbial cell 'ghosts" at Yaman-Kasy. A comparison between the known fossil assemblages and grouped modern vent communities (Fig. 6) shows that all but one of the in external moulds. taxonomic groups at this level of analysis are The identification of the fossil vent tubes is shared. more problematic because many modern marine No arthropod fossils have been found at fossil taxa produce tubes with the potential for vent sites. This may be because arthropods did fossilization and intraspecific variation in tube not occur in ancient vent communities, or morphology is common (e.g. the vent vestimen- because they did occur but were not fossilized. tiferan Ridgeia piscesae, Southward et al. 1995). The presence of crustaceans in a Carboniferous Identification of fossil tubes from vents (Hay- hydrocarbon seep assemblage (von Bitter et al. m o n & Koski 1985; Kuznetsov et al. 1991a, 1992; R. Fortey pers. comm.) shows that 1993; Little et al. 1997) and SEDEX deposits arthropods have lived in chemosynthesis-based (Banks 1985; Moore et al. 1986; von Bitter et al. communities in the past and this may lend 1992; Russell 1996) has been made using modern weight to the second explanation. tubicolous vent animals (alvinellid polychaetes Little et al. (1997) stated that monoplacoand vestimentiferans) for comparison, the in- phorans and lingulate brachiopods (representing ference being that fossil vent communities had a class and a phylum, respectively) as vent taxa similar taxonomic elements to modern ones. were unique to fossil assemblages. However, However, there is danger of circularity in this monoplacophorans (as yet unidentified) have argument and the possibility remains that the recently been recovered on the shells of Bathfossil vent tubes were made by unknown, ymodiohts at the Menez Gwen vent site, Midperhaps extinct, tubicolous vent organisms. Atlantic Ridge, 37°50'N (A. War6n, pers. For this reason the identification of the fossil comm.). There are also two records of brachiovent tubes as vestimentiferans and alvinellid pods from mid-ocean ridges: terebratulids have polychaetes is tentative. been dredged from fresh lava flows on the Southeast indian Ridge (Gregory & Lee 1995); and unidentified brachiopods have been colC o m p a r i s o n s b e t w e e n j o s s i l vent a s s e m lected from inactive sulphide mounds at the Escanaba Trough (Van Dover et al. 1990). At blages a n d m o d e r n vent c o m m u n i t i e s present it is not clear whether the Escanaba The fossil vent assemblages are taxonomically Trough brachiopods are utilizing low temperaheterogeneous containing brachiopods (lingu- ture hydrothermal fluids or are simply elements c

"5 "~ .~_ c

268

C. T. S. LITTLE ET AL.

of a normal hard-substrate, non-vent c o m m u nity. All the fossil shelly v e n t taxa are extinct a n d p r e l i m i n a r y t a x o n o m i c analysis suggests n o n e o f these are ancestors o f any m o d e r n vent molluscs or ocean-ridge b r a c h i o p o d s . It follows t h a t the m o d e r n vent e n v i r o n m e n t is n o t a refuge for the k n o w n P a l a e o z o i c a n d M e s o z o i c shelly vent taxa a n d there has been m o v e m e n t o f t a x o n o m i c g r o u p s in a n d o u t o f the vent e c o s y s t e m t h r o u g h time. T h e n u m b e r o f shared higher g r o u p s between fossil vent assemblages a n d m o d e r n vent c o m m u n i t i e s w o u l d seem to s u p p o r t N e w m a n ' s (1985) refuge h y p o t h e s i s for vent c o m m u n i t i e s . H o w e v e r , t w o o f these s h a r e d g r o u p s , the vestimentiferans a n d the alvinellid polychaetes, c a n n o t be used in defence o f either this or H i c k m a n ' s (1984) r a p i d a d a p t i v e r a d i a t i o n / extinction hypothesis. T h e definition o f a refuge is a place that animals have retreated into or been left in d u r i n g a p e r i o d o f adversity. On available evidence, vestimentiferans a n d alvinellids have only ever lived at h y d r o c a r b o n seeps a n d / o r vent sites. These habitats are therefore n o t refuges for them. T h e p a t t e r n s h o w n by the fossil v e n t m o l l u s c s a n d b r a c h i o p o d s m a y s u p p o r t H i c k m a n ' s (1984) h y p o t h e s i s but a m a j o r p r o v i s o is the small n u m b e r o f taxa involved (seven) a n d that m u c h t a x o n o m i c w o r k r e m a i n s to be done. We suggest that the fossil evidence m a y s u p p o r t the finding by M c A r t h u r & Tunnicliffe (this volume), based on m o l e c u l a r and morphological phylogenetic work on modern vent taxa, t h a t there is a range o f lineage ages present at vent sites. We wish to thank all the Russian mining kombinats that allowed access to the relevant open pits in the Urals, in particular the M M C K and CMK companies. CTSL thanks A. War~n and V. Tunnicliffe for unpublished information about modern vent communities, J. Cann for his foresight about fossil sites in Cyprus and help in the field, the Cyprus Geological Survey, and C. Van Dover and two anonymous reviewers for criticism of an earlier draft of this paper. This research is supported by N ERC grant GR3/10903 and a Royal Society joint project with the FSU.

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This volume. McLEAN, J. H. & HASZPRUNAR,G. 1987. Pyropeltidae, a new family of cocculiniform limpets from hydrothermal vents. Veliger, 30, 196 205. MACLEOD, C. J. & ROTHERY, D. A. 1992. Ridge axial segmentation in the Oman ophiolite: evidence from along-strike variations in the sheeted dyke complex. In: PARSON, L. M., MURTON, B, J. & BROWNING, P. (eds) Ophiolites and their Modern Oceanic Analogues. Geological Society, London, Special Publications, 60, 39-63. MASLENNIKOV, V. V. 1991. The Lithological Control of

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VON BITTER, P. H., SCOTT, S. D. & SCHENK, P. E. 1992. Chemosynthesis: an alternative hypothesis for Carboniferous biotas in bryozoan/microbial mounds, Newfoundland, Canada. Palaios, 7, 466~484. WALTER, M. R. 1996. Ancient hydrothermal ecosystems on Earth: a new palaeobiological frontier. In: BOCK, G. R. & GOODE, J. A. (eds) Evohttion of hydrothermal ecosystems on Earth (and Mars?). Wiley, Chichester, 112-127. WAREN, A. & BOUCHET, P. 1986. Four new species of Provanna Dall (Prosobranchia, Cerithiacea?) from the East Pacific hydrothermal sites. Zoologica Scripta, 15, 157 164. YAZEVA, R. G., MOLOSHAG, V. P. & BOCHKARYOV,V. V. 1992. Geology oJ" the Sqfyanovka sulphide deposit (the MMdle Urals). Urals Branch of the Russian Academy of Sciences, Ekaterinburg. ZAYKOV,V. V., MASEFNNtKOV,V. V,, ZAYKOVA,E. V. & HERRINGTON, R. J. 1996. Hydrothermal activity and segmentation in the Magnitogorsk-West Mugodjarian zone on the margins of the Urals palaeo-ocean. In: MACLEOD,C. J., TYLER, P. A. & WALKER, C. L. (eds) Tectonic, Magmatic, Hydrothermal and Biological Segmentation of MidOcean Ridges. Geological Society, London, Special Publications, 118, 199-210. ZONENSHAIN, L. P., KORINEVSKI, V. O., KAZMIN,V. O., PECHERSKI, D. M., KHA1N, V. V. & MATEENKOV, V. V. 1984. Plate tectonic model of the southern Urals development. Tectonophysics, 109, 95-135.

Relics and antiquity revisited in the modern vent fauna A. G. M c A R T H U R 1 & V. T U N N I C L I F F E

Department o f Biology, University of Victoria, Victoria, B.C., V8 W 2 Y2, Canada 1Present address. Marine Biological Laboratory, 7 M B L Street, Woods Hole, M A 02543-1015, USA

Abstract: A total of 464 species is currently known from hydrothermal vents, with 82% being endemic. More strikingly, endemicity at hydrothermal vents reaches high taxonomic levels with 18 families, 4 superfamilies, 2 suborders, and 1 order endemic. It has been proposed that this magnitude of endemicity is reflective of an ancient, refugial fauna. A review of the many endemic groups reveals that hydrothermal vent communities have mosaic origins. Many endemic groups have recent origins while some appear to reflect ancient faunal elements that have survived in hydrothermal vent refugia. Some may be members of groups surviving refugially throughout the deep-sea while others may have survived refugially in a variety of sulphide-rich habitats, such as hydrothermal vents, hydrocarbon seeps and sunken whale bones. Evolution of sulphide tolerance and especially symbioses with chemoautotrophic bacteria may be threshold adaptations for invasion of multiple sulphide-rich environments. These environments may provide protection from certain causes of extinction. In addition to reviewing current views on endemic taxa, a molecular systematic investigation of the evolutionary origins of vent endemic neomphalinid gastropods is presented.

The concept of a 'lost world' is a fascinating one, not only for historians but also for biologists. We are tantalized by discoveries of living fossils such as Neopilina the deep-sea monoplacophoran, Latimeria the coelacanth, and Metasequoia the dawn redwood. If individual hold-overs from the past can remain in the modern world, perhaps whole qost' faunas and floras remain to be discovered. U n d e r marginal or isolated conditions a habitat may remain relatively immune from invasion and displacement by superior competitors. Perhaps the archetypal example of such a process is the Australian continent where entry of marsupial animals coincided with the breakup of Gondwanaland and glaciation of the nearest landmass-Antarctica. Other examples of island refuges for early Cenozoic organisms include certain floras of New Zealand and the primates of Madagascar. Isolation in an island situation seems to be a crucial factor. In the aquatic reahn, isolated habitats such as old lakes (e.g. Baikal, several African lakes; Martens 1997) as well as submarine and anchihaline caves provide other possibilities for isolation (Danielopol 1990; Kase & Hayami 1992; Jaume & Boxshall 1996: Sarbu et al. 1996). The term 'relic' is applied loosely. Zezina (1994) points out that the term has been in the literature for 135 years and includes a delayed

rate of evolution in a group formerly abundant but now relegated to marginal areas by more advanced organisms. Just how old a survivor of an otherwise extinguished lineage has to be to attain 'relic' status is not clear.

Antiquity in the deep-sea There remains an on-going debate on whether elements of the deep-sea fauna represent holdovers from Mesozoic and even upper Palaeozoic eras. In general, evolutionary novelties corresponding to new supeffamilies or higher taxa do not seem to emerge in deep-sea habitats, but rather shallow water and shelf habitats are the sources of major taxon originations that gradually propagate down-slope (Jablonski et al. 1983; Sepkoski & Miller 1985). Given the vast size of the h a b i t a t - t h e bathyal and abyssal zones occupy about 60% of the planet the possibility that some ancient inhabitants still exist remains intriguing. Zezina (1994) argues that the bathyal zone (500 to 3000m) has remained the most stable marine realm, escaping effects of sea-level changes above and the influences of anoxic incursions at greater depths. In the 1950s and 1960s, recovery of new organisms phylogenetically allied with older lineages provoked considerable argument about the nature of the deep ocean fauna (Ekman

MCARTHUR, A. G. & TUNNICLIFFE,V. 1998. Relics and antiquity revisited in the modern vent fauna. In: MILLS, R. A. & HARRISON,K. (eds) Modern Ocean Floor Processes and the GeologicalReeord, Geological Society, London, Special Publications, 148, 271 291

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Table 1. Endemicity at h)'drothermal vents. Fish and cephah)pod records have been excluded as their mobility and vagrant nature make endemicity impossible to assess. Percentage information on subfamilies, supelJhmilies, suborders, and subclasses is excluded as these rankings are not used consistently throughout the phyla examined. Endemicity of animals found in hydrothermal vents' is presented for four Jaunal groups, each more inclusive than the previous." hydrothermal vent communities alone; hydrothermal vent and hydrocarbon seep communities combined," vent, seep, and communities on deep-sea sunken bone combined," vent, seep, and communities on deep-sea sunken bone, wood and algae combined. Animal taxa not endemic" to these four groupings are found in other marine habitats and may or may not be restricted to the &,ep-sea. Association with hydrocarbon seeps has' been assumedJor some reports o f vesthnent(feran pogonophorans (e.g. Webb 1969)

Taxon Species Genus Subfamily Family Superfamily Suborder Order Subclass Class Phylum

Total number in vents

Vent endemic

Vent and seep endemic

Vent, seep, and bone endemic

464 240 n/a 119 n/'a n/a 51 n/a 20 9

381 (82.1%) 107 (44.6%) 5 18 (15.1%) 4 2 I (2.0%)

+ 12 (84.7%) +9 (48.3%) +2 +4 (18.5%) +1 +1 + 1 (4.0%) +1

+4 (85.6%) +2 (49.2%) +1 + 1 (19.3%) -

Vent, seep, bone, and wood endemic +6 (51.7%) + 1 (20.2%)

-

-

1953; Menzies & lmbrie 1955; Zenkevich & Birshtein 1960; Madsen 1961; among others). The controversy retreated in the face of speculation that major shifts in deep ocean temperatures could have caused widespread extinctions of deep-water faunas (i.e. Bruun 1956; Menzies et al. 1973). Nonetheless, the antiquity concept remains alive due to a continuous trickle of new finds of animals in families or superfamilies known only as Mesozoic or early Tertiary fossils; groups include crinoids, bivalves, gastropods, crustaceans, bryozoans and brachiopods (see Zezina 1994). That these organisms represent minor anomalies in an otherwise 'youthful' deep-sea fauna remains a bone of contention (Kuznetsov 1991). As evidence accumulates of widespread anoxia in deep waters during parts of the Mesozoic and Cenozoic (i.e. Kennett & Stott 1991: Wignall & Twitchett 1996; Isozaki 1997) the survival of even remnants of ancient faunas becomes all the more fascinating.

Antiquity at hydrothermal vents The concept of antiquity at deep-sea vents arose from the observation that nearly all animals collected at vents were undescribed species and that many of these species appeared to have 'antiquated' or plesiomorphic features. In 1979, Newman described a barnacle from East Pacific Rise vents as a Mesozoic relic because of primitive features akin to species known from Jurassic deposits (Newman 1979). Shortly afterwards, McLean (1981) described a common vent

limpet as a remnant from a gastropod lineage that radiated in the Palaeozoic but declined in the Mesozoic. Several observations of this ilk moved Newman (I985) to entitle his paper: 'The abyssal hydrothermal vent invertebrate fauna: a glimpse of antiquity?' At that time, 58 novel species were known from vents and many of these were also assigned to new genera and new families. Newman's observation was that, if the majority of genera (and many families) at vents were new to science, they must have been sequestered in this distinct habitat for a long time. For example, deep-sea barnacles with a fossil record have genera of ages between 45 and 162Ma ( N e w m a n 1985) and most marine gastropod families have a fossil record back to the late Cretaceous (Tracey et al. 1993). To generate new higher taxa that have diverged from known barnacle genera and gastropod families would take even longer. Thus it was the combination of plesiomorphic characters in some vent animals, the high degree of endemism at generic and familial levels, and the corroboration with fossils in a few cases that provoked the antiquity concept. Tunnicliffe (1992) re-examined the available information seven years later when the known vent species numbered 236. Endemism was then 95% at the species level, 58% at the generic level, and 22% at the familial level, much higher than any other marine habitat. At that time, information from seep and whale bone habitats began to appear. While few of these species are shared with vents (Tunnicliffe et al. 1996), there are clear relations

ANTIQUITY IN THE MODERN VENT FAUNA

273

Table 2. Higher taxa endemic to hydrothermal vents, with notes on endemiciO, associations with communities of hydrocarbon seeps, sunken bone, wood, and algae. * represents a resurrection of taxa otherwise only known Ji*om the fossil record (see text)

Taxon Subclasses Pogonophora Obturata Jones 1981 Orders Pogonophora Axonobranchia Jones 1985 Basibranchia Jones 1985 Suborders Gastropoda Euomphalina* McLean 1981 Lepetopsina* McLean 1990 Crustacea Brachylepadomorpha* Withers 1923 Superfamilies Gastropoda Lepetodriloidea McLean 1988 Lepetopsoidea* McLean 1990 Neomphaloidea McLean 1981 Peltospiroidea McLean 1989 Crustacea Bythograeoidea Williams 1980 Families Pogonophora Riftiidae Jones 1981 Alaysiidae Southward 1991 Arcovestiidae Southward & Galkin 1997 Lamellibranchiidae Webb 1969 Ridgeiidae Jones 1985 Tevniidae Jones 1985 Gastropoda Lepetodrilidae McLean 1988 Gorgoleptidae McLean 1988 Clypeosectidae McLean 1989 Cyathenniidae McLean 1990 Neomphalidae McLean 1981 Peltospiridae McLean 1989 Neolepetopsidae McLean 1990 Provannidae Waren & Ponder 1991 Pyropeltidae McLean 1987 Crustacea Alvinocarididae Christoffersen 1986 Bythograeidae Williams 1980 Ecbathyriontidae Humes 1987 Dirivultidae; Humes & Dojiri 1980 Brachylepadidae* Woodward 1901 Neoverrucidae Newman & Hessler 1989 Poh,chaeta Alvinellidae Desbruyeres & Laubier 1986 Archinomidae Kudenov 1991 Hemichordata Saxipendiidae Woodwick & Sensenbaugh 1985 Subfamilies Gastropoda Sutilizoninae McLean 1989 Temnocinclinae McLean 1989 Bivah, ia Bathymodiolinae Kenk & Wilson 1985 Crustacea Eolepadinae* Buckeridge 1983 Poh, chaeta Branchinotogluminae Pettibone 1985 Branchiplicatinae Pettibone 1985 Branchipolynoinae Pettibone 1984 Lepidonotopodinae Pettibone 1983

Affinity

Habitats

= Vestimentifera

vents, seeps

Obturata Obturata

vents vents, seeps

uncertain Patellogastropoda

vents vents, seeps

Sessilia

vents

Vetigastropoda Lepetopsina Euomphalina Euomphalina

vents vents, seeps vents vents

Brachyura

vents

Axonobranchia Basibranchia Basibranchia Basibranchia Basibranchia Basibranchia

vents vents vents vents, seeps vents vents

Lepetodriloidea Lepetodriloidea Fissurelloidea Neomphaloidea Neomphaloidea Peltospiroidea Lepetopsoidea Loxonematoidea Lepetellida

vents vents vents vents vents vents vents, seeps vents, seeps, sunken wood vents, seeps, whale bones

Decapoda Bythograeoidea Siphonostomatoida Siphonostomatoida Brachylepadomorpha Verrucomorpha

vents, seeps vents vents vents, seeps vents vents

Phyllodocida Phyllodocida

vents vents

Enteropneusta

vents

Scissurellidae Scissurellidae

vents vents

Mytilidae

vents, seeps

Scalpellidae

vents

Polynoidae Polynoidae Polynoidae Polynoidae

vents, whale bones vents vents, seeps vents

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A.G. McARTHUR & V. TUNNICLIFFE

Table 3. Taxonomic analysis of endemici O' at hydrothermal vents. Values represent number of taxa f o u n d at hydrothermal vents endemic to vents alone or endemic to a combination of hydrothermal vents, hydrocarbon seeps, sunken bone or wood. Fish and cephalopod records have been excluded as in Table 1. Data are presented as." number of taxa found at vents that are endemic to the above sulphide-rich environments / total number of taxa found at vents (percentage o f taxa found at vents endemic to sulphide-rich environments). For example, eight famgies o.1" pogonophorans are known at hydrothermal vents, six of which (i.e. 75%) are endemic to sulphide-rich environments such as hydrothermal vents, hydrocarbon seeps, sunken bone or wood Taxon Pogonophora Polychaeta Gastropoda Bivalvia Crustacea Cnidaria Others

Family

Genera

Species

6/8 (75.0%)

7/9 (77.8%) 21/56 (37.5%) 46/60 (76.7%) 2/8 (25.0%) 39/76 (51.3%) 2/11 (18.2%) 4/7 (57.1%)

1t/11 (100%) 89/104 (85.6%) 113/130 (86.9%) 20/23 (87.0%) 131/152 (86.2%) 7,/12 (58.3%) 26/30 (86.7%)

2/20 (10.0%)

9/26 (34.6%) 0/5 (0%) 6/41 (14.6%) 0/6 (0%) 1/13 (7.7%.)

at higher taxonomic levels. Information from fossil sites was tantalizing but very sparse (Tunnicliffe 1992) and did not clearly define affinities for any taxa. With each publication of new vent taxa, the antiquity hypothesis is tested further. Recent work on fossil vents has called into question the validity of the concept (Little et al. 1997). In this paper, we review the status of the vent fauna from the points of view of endemism and apparent 'antiquated' form. We also present one test of the hypothesis which explores the origins of vent gastropod groups using molecular techniques.

Endemism at hydrothermal vents While most vent species are new to science, with on-going exploration of non-hydrothermal environments we can anticipate the continued discovery of 'vent species' not associated with hydrothermal vents. Nonetheless, endemicity remains extremely high (Table 1). It is notable that exploration beyond the initially explored eastern Pacific vents to mid-Atlantic and western Pacific ridges and back-arc basins has continued to yield new endemic species and higher taxa. Globally, high endemism appears to be the standard characteristic of hydrothermal vent communities. Tunnicliffe et al. (1996) postulated the existence of a deep-sea sulphophilic fauna reflective of long-term in situ evolution at sulphide-rich deep-sea environments. The deepsea does appear to host an endemic sulphophilic fauna known from hydrothermal vents, hydrocarbon seeps, sunken vertebrate bones, and submerged wood (Table 1). Endemic genera, families, and higher taxa may possibly reflect long-term in situ radiation and evolutionary

association with these sulphide-rich habitats. Endemic higher taxa are presented in Table 2. Endemic elements are known from the Pogonophora, Polychaeta, Crustacea, Gastropoda, Bivalvia and Hemichordata (Table 2). While species of the Gastropoda, Polychaeta, and Crustacea numerically dominate hydrothermal vents and species-level endemicity is high for all phyla known at vents except the Cnidaria, distribution of higher-level endemicity is uneven (Table 3). The Pogonophora and Gastropoda exhibit the highest family and genus-level endemicity while endemicity for the Polychaeta and Bivalvia is more predominant for species and genera. Endemicity for the Crustacea is most predominant for genera but six families are endemic: two Decapoda, two Cirripedia, and two Copepoda. Taxonomic levels for the Pogonophora will shift dramatically if the group is demoted to familial status as recommended by McHugh (1997) and Rouse & Fauchald (1997). The two major explanations presented by Newman (1985) for the vent endemicity levels were long-term in situ radiation or rapid morphological change after entry into the habitat, in the majority of cases, endemic elements in the hydrothermal vent fauna represent independent invasions of sulphide-rich habitats as endemic families (or genera) within phyla and are known from different orders, suborders, or superfamilies. The next section reviews what is known about the evolutionary age and origins of the various endemic groups.

Review of the major endemic groups Pogonophora

The systematic placement of the vent tubeworms, 'vestimentiferans', has a convoluted

ANTIQUITY IN THE MODERN VENT FAUNA history and a resolution that is currently unclear. Webb (1969) placed the first vestimentiferan species (dredged most likely from seeps near Oregon) in a new order of the phylum Pogonophora recently erected by Ivanov (1963). After the discovery of vent tubeworms, Jones (1985) raised vestimentiferans to the rank of phylum after previously introducing the name Obturata for them (Jones 1981). The phylum status has not met with general acceptance (i.e. Southward 1988; Southward & Galkin 1997). There exists a legacy of opinion that both pogonophorans and vestimentiferans are closely allied to, if not a sub-taxon of, the Annelida (van der Land & Norrevang 1975). Recent work on morphological characters finds close similarities to annelids (Bartolomaeus 1995; Rouse & Fauchald 1997; but see Malakhov et al. 1996). Molecular investigations of haemoglobins (Suzuki et al. 1993) and elongation factor-lo~ (Kojima et al. 1993; McHugh 1997) support close relation to polychaete annelids. Small subunit ribosomal R N A sequences first favoured a common ancestor with the Echiura and close relation to the Mollusca, not the Annelida (Winnepenninckx et al. 1995), but with inclusion of more molluscan sequences, they favoured an unresolved Annelida-Echiura-Pogonophora clade (Winnepenninckx et al. 1996). Aguinaldo et at. (1997) present a discussion of analytical problems with small subunit ribosomal R N A sequences when examining metazoan phylogeny. Much of the difficulty in resolving phylogenetic relations lies in availability, firstly, of a complete set of morphologic characters and, secondly, a representative suite of tubeworms and appropriate outgroups. For this study, we use the interim taxonomy of Southward & Galkin (1997) and recognize that these animals can be placed either in the phylum Annelida or in the phylum Brachiata, 'depending on the future conclusions about the embryonic origin of the mesoderm.' (p. 485, Southward 1988). Recent embryological and larval comparisons suggest that, unlike perviate pogonophorans, enterocoelous mesoderm formation will not be found in the vestimentiferans and that they may indeed be specialized annelids (Young et al. 1996). McHugh's (1997) recent investigation using elongation factor-lo~ D N A sequences strongly favoured polychaete affinities of vestimentiferan pogonophorans (RMgeia, Lamellibrachia) and McHugh (1997) recommended resumption of the annelid family taxon Siboglinidae for the Pogonophora. In contrast, extensive examination of the microscopic anatomy of the vestimentiferan R i d g e i a p h a e o p h i a l e led Malakhov et al. (1996) to propose close relation

275

of the Pogonophora to the Chordata. At the moment, the question of antiquity and 'relic' status of the vestimentiferans must await the result of ongoing phylogenetic studies. However, ancient sulphide deposits of both Palaeozoic and Mesozoic age contain organic remains with many similarities to modern vent and seep tubeworms (Haymon et al. 1984; Little et al. 1997; Little et al. this volume). The vestimentiferans may be a long-term vent resident, even though modern representatives exhibit low molecular divergence (Suzuki et al. 1993; Feldman et al. 1997). Feldman et al. (1997) present molecular evidence that modern vestimentiferan lineages may have diverged only 10Ma ago, suggestive of historical crises or faunal change for the group during the Cenozoic. Polychaeta

Pettibone (1983, 1984, 1985a, b) has described four new subfamilies of polynoid polychaetes from hydrothermal vents, two of which are now known from other sulphide-rich environments (Table 2). Each subfamily was recognized on the basis of morphological novelties, particularly well-developed branchiae. Their evolutionary age and length of association with sulphide-rich environments is unknown. It is additionally unknown if they represent independent invasions. The Alvinellidae represents a large endemic family containing two genera and eleven species known from both eastern and western Pacific vents. The worms live in tubes and are located very close to hydrothermal emissions. They experience extremes in temperature, dissolved oxygen, heavy metal precipitation, natural radioactivity, and dissolved hydrogen sulphide (reviewed by Gaill & Hunt 1991). Desbruy~res & Laubier (1991) hypothesized ancient origins of the family at the time when the ampharetid and trichobranchid lineages of the Terebellomorpha appeared. The Alvinellidae exhibits plesiomorphies for the Terebetlomorpha unknown from other living families, such as the presence of nodopodia throughout the body and differentiation between thorax and abdomen. Cladistic analysis of the Polychaeta supports the monophyly and ampharetid alliance of the family (K. Fauchald, pers. comm.). Feral et al. (1994) used partial large-subunit ribosomal RNA sequences to examine the phylogenetic position of the Alvinellidae within the Terebellomorpha. They found the Alvinellidae to be a monophyletic sister clade to the Terebellidae and support the phylogenetic novelty of the group within the

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Terebellomorpha; the family does not appear to be a derivative of other deep-sea terebellimorph families. Fossil tubes in ophiolite sulphide deposits from the Palaeozoic to Cenozoic (e.g. Oudin et al. 1985; Boirat & Fouquet 1986; Kuznetsov et al. 1990; Little et al. 1997) have been ascribed to alvinellids but there are too few characters to be sure. A new polychaete species Euphrosine rosacea Blake 1985 has been reassigned to a new genus, Archinome and a new family, Archinomidae (Kudenov 1991). The author recognizes confusion within the original family (Amphinomidae) and suggests a common ancestry. The parent order, Amphinomida, "represents an ancient, monophyletic, and strongly isolated order of polychaetes ... Archinome probably reflects a long term association with hydrothermal habitats and isolation from other Amphinomida.' (Kudenov 1991, p. 118). Fossil Amphinomida are known from the Carboniferous (Briggs & Kear 1993). Thus, in both the Alvinellidae and the Archinomidae, we find vent-endemic families with many characters plesiomorphic for their respective orders that may reflect long-term isolation. The family Nautiliniellidae (previously Nautilinidae), was originally described from hydrothermal vents and hydrocarbon seeps (Miura & Laubier 1989; Blake 1990; Miura & Laubier 1990), but Blake (1993) described new species and genera elsewhere. Flascarpia alvinae is likely from a hydrocarbon seep environment while the environment of Santalema miraseta is uncertain. A third genus, represented by Miura spinosa, is located in a deep-sea sediment environment of low oxygen concentration. In a cladistic examination of the morphological features of the Nereidoidea, Glasby (1993) found that only the autapomorphy of absent anal cirri defined the Nautiliniellidae as separate from other nereidoid families and that the family does not root deeply in the nereidoid phylogenetic tree, but instead probably shares a (recent?) common ancestor with the Pilargidae or Syllidae. It is possible that this family, and many vent-associated polychaete genera, may belong to groups that are simply tolerant of low concentrations of dissolved oxygen and does not reflect long-term isolation at hydrothermal vents. Crustacea

Of the 63 species of copepods at vents and seeps, 47 belong to the vent and seep endemic family Dirivultidae (Siphonostomatoida). This family of copepods has been found parasitic on vestimentiferan tube worms, alvinellid polychaetes, and various crustaceans, as well as

living independently (e.g. Humes 1987; Humes & Lutz 1994). The family contains one genus and species endemic to hydrocarbon seeps (Dirivultus dentaneus; Humes & Dojiri 1980); the rest are vent-specific from all known vent locations. Huys & Boxshall (1991) postulate that the Dirivultidae represents the earliest branch of siphonostome copepod evolution, based on an abundance of plesiomorphic features. They suggest that the deep-sea (presumably meaning hydrothermal vents) may have acted as a refuge for this early lineage. Similarly, while they do not place the Ecbathyriontidae, a hydrothermal vent endemic copepod family represented only by Ecbathyrion prolixicauda (Humes 1987), within phylogenetic context, they suggest this plesiomorphic family has found the same refuge. Within the poecilostomatoid copepods, the family Erebonasteridae, formerly endemic to hydrothermal vents, is now known from hydrocarbon seeps, non-sulphide deep-sea, and the continental shelf (Humes 1987; Humes 1989; Huys & Boxshall 1990; Huys 1991). Huys & Boxshall (1991) consider the family to be the first branch of the poecilostomatoid radiation and one of the species is considered the most primitive poecilostomatoid copepod known (Huys & Boxshall 1990). As more species are discovered, examinations of the phylogeographic history of the family would be of interest, particularly in terms of the bathymetric origins of these modern representatives. In overall copepod phylogeny, Ho (1990) and Huys & Boxshall (1991) place the Siphonostomatoida and Poecilostomatoida (three endemic vent genera) as the most derived of the copepod orders, the major distinguishing characters being adaptations to a parasitic existence. The one described vent calanoid is a new genus of Spinocalanidae; while the genus is derived, the family is the most primitive within its superorder (Fleminger 1983). A recently discovered poecilostome from the Juan de Fuca Ridge displays several characters that indicate a very primitive form (A. Humes pers. comm.). Four genera of scavenging and carnivorous decapods found in hydrothermal vents have been placed in the endemic superfamily Bythograeoidea (family Bythograeidae). Like the endemic copepods, the bythograeids are known from hydrothermal locations of the Atlantic and both sides of the Pacific (Hessler & Martin 1989). It is unclear if the superfamilial ranking of the Bythograeoidea is reflective of long-term association or rapid, post-invasion morphological change and diversification. While brachyuran superfamilies have origins in late Mesozoic, A. B. Williams (cited in Guinot & Segonzac 1997)

ANTIQUITY IN THE MODERN VENT FAUNA

277

Acrothoracica Ascothoracida CYPRILEPAS Heteralepadidae Iblidae

Outgroups

PRAELEPAS Lepas

ARCHAEOLEPAS Neolepas *

r

-BRACH~'LEPADOMORPHA Neoverruca *

Capitulum Pollicipes Scalpellum Scillaelepas Verrucidae PROVERRUCA EOVERRUCA Chionelasmus

Scalpelloidea Verrucomorpha

Eochionelasmus *

Catophragmus Chthamalinae Pachylasma

Balanomorpha

WAIKALASMA Octomeris

Fig. 1. A reproduction of Glenner et al.'s (1995) phylogenetic hypothesis for the cirriped barnacles, representing a strict consensus of the most parsimonious trees based upon anatomical features. Extinct taxa are presented in capitals and hydrothermal vent endemic taxa, considered the most primitive representatives of their suborders, are marked with an asterisk. The phylogenetic position and Mesozoic origins of NeoIepas and Eochionelasmus are supported, although Neolepas was not monophyletic with the other Scalpellomorpha. Resolution among the cirriped suborders was not obtained. The unexpected placements of Neoverruca and Capitulum were not strongly supported. Since publication of this hypothesis, the only known living brachylepadomorphan has been discovered at hydrothermal vents (Neobrachylepas, Newman & Yamaguchi, 1995) and living Waikalasma has been discovered at abyssal depths off the New Hebrides (Buckeridge 1996).

believes the Bythograeoidea to have diversified in the Cenozoic. Christoffersen (1986, 1991) placed the vent and seep endemic decapod shrimp genera A lvinocaris and Rimicaris in the new family Alvinocarididae, upon a cladistic examination of the Caridea. Segonzac et al. (1993) redefined the family to include Chorocaris, but Martin & Christiansen (1995) did not support the new family and favoured retention in the Bresiliidae despite its likely artificial nature. Similarly, T. Shank finds no justification for the novel family and that molecular studies show that extant bresiliid species (including those of Alvinocarididae) represent recent radiation (T. Shank, in preparation). For our purposes, we include the Alvinocarididae in our listing of endemic families but note that future discoveries and research may dissolve the family. A solid argument for antiquity at hydrothermal vents comes from the barnacles. There are

currently seven genera of barnacles known from hydrothermal vents, all but two of which are endemic at higher levels. Of four suborders of the Thoracica (attached, sessile barnacles), the most primitive living representative of the Scalpellomorpha (i.e. Neolepas; Newman 1979), Verrucomorpha (i.e. Neoverruca; Newman & Hessler 1989), and Balanomorpha (i.e. Eochionelasmus; Yamaguchi & Newman 1990) as well as the only living representative of the Brachylepadomorpha (i.e. Neobrachylepas; Newman & Yamaguchi 1995) are only known from hydrothermal vents. Not only are they plesiomorphic for their suborders, but they have clear affinities to barnacles known from the fossil record of the Mesozoic, particularly the Jurassic (Newman 1979; Buckeridge & Grant-Mackie 1985; Newman & Hessler 1989; Newman & Yamaguchi 1995). A fossil Neolepas is known from a Lower Jurassic continental shelf environment, indicating restriction to hydrothermal vents was during

278

A. G. McARTHUR & V. TUNNICLIFFE

FHeterobranchiaI

Mesodon

I NeritimorphaI Littorina& ~ Buccinum~ ~o Ampullaria~ ".... ." Campa~% ..°'. Viviparus ' ...... ~

[ Caenogastropoda]

i"

Temn°cinclis~98% DioOora, ~

Nerita& Olgasolaris T~heoduxos athynerita& Shinkailepas

~

Rhynchopelta Depressigyra "~CYM:tl~:r°dT&ta Peltospira

M.°lnn°cdTnsta&~-'~~j ~ ( 6 X

I 'Neomphalina'I

~'~e uamymargames

[ Vetigastropoda[ Fig. 2. 28S rRNA/rDNA investigation of the evolutionary origins of the living, hydrothermal vent endemic Euomphalina (i.e. Neomphalina). After exclusion of four hyper-variable subdomains, evaluation of the most parsimonious trees revealed sequence evolution followed a Kimura's two-parameter model (i.e. transition bias) with some sequence positions invariant and the others evolving at rates differing according to a gamma distribution. As these phenomena violate assumptions of parsimony, all analyses were performed using maximum likelihood under the above K2P+ I + F model with parameters estimated from the data. The tree presented represents the unrooted ingroup topology as outgroup sequences proved to be randomized relative to the ingroup (i.e. overly distant). The Patellogastropoda sequences were excluded as they exhibited a systematic two-fold increase in evolutionary rate and thus were additionally randomized relative to the ingroup, as previously noted by Tillier et al. (1994). The dotted line represents the portion of the tree not supported by bootstrapping of parsimony, distance, and likelihood methods and is equivalent to radiations of the early to mid-Palaeozoic. Percentages represent bootstrap support of distance methods under the K2P+ I + F model for the supported monophyletic gastropod suborders (in solid lines: Heterobranchia, Caenogastropoda, Vetigastropoda, Neritimorpha, Neomphalina). Similar support was found for parsimony and preliminary maximum likelihood bootstrapping. See the text for a discussion of the implications. Discussion of phylogenetic patterns and support within the major groups will be presented in a future publication.

or after the Jurassic (Buckeridge & GrantMackie 1985, but see Newman & Yamaguchi 1995). Neobrachylepas relica represents the first living Brachylepadomorpha, a suborder of Jurassic origins, since their extinction in the Miocene (Newman & Yamaguchi 1995). Glenner et al. (1995) performed a cladistic examination of the morphology of both living and extinct thoracican barnacles, including the 'primitive' vent endemic Neolepas, Neoverruca, and Eochionelasmus. Their phylogenetic hypothesis is reproduced in Fig. 1. While their phylogenetic tree differed from aspects of noncladistic interpretations, it did support the suggested phylogenetic novelty of the ventendemic barnacles and thus their Mesozoic affinities, although the exact affinities of Neoverruca were uncertain (Brachylepadomorpha

or Verrucomorpha?). Glenner et al. (1995) hypothesize C a r b o n i f e r o u s origins for the Scalpellomorpha and Neolepas based on their interpretation of Pabulum spathiforme as a scalpellomorph, a contention not supported by W. A. Newman (pets. comm.) who favours Triassic origins (e.g. Newman & Yamaguchi 1995). Hydrothermal vents aside, the deep-sea in general harbours many relic barnacles as exhibited by the rediscovery of the most primitive living eight-plated balanomorph, Waikalasma, previously only k n o w n from the Miocene (Buckeridge & Newman 1992), at bathyal depths off the New Hebrides (Buckeridge 1996). Many older barnacle genera, often with relic Tethyan distributions, have been displaced to the deepsea by more recent forms (Stanley & Newman 1980; Newman & Foster 1987). One of these

ANTIQUITY IN THE MODERN VENT FAUNA displaced genera, Scillaelepas, is also known at hydrothermal vents. Vents sequester the most primitive forms of four lineages of barnacles. The habitat appears to offer these animals some benefit. Yamaguchi (1994) notes that a convergent adaptation is present in the vent forms: the presence of finely setose mouthparts capable of extracting fine suspended particles like bacteria. A new Neolepas from Indian Ocean vents shows the same feature (Southward et al. 1997). Gastropoda The most extensive and diverse endemicity at hydrothermal vents is found in the Gastropoda; two subfamilies, nine families, four superfamilies, and two suborders are restricted to vents or a combination of vents and other sulphide-rich habitats. While gastropods have an excellent marine fossil record, it is poor for the deep-sea and non-existent for vents. Antiquity for most vent gastropods is postulated on the basis of plesiomorphic or novel anatomy, with some tentative commentary on fossil associations. We summarize here what is known and include some results of a molecular systematic investigation by one of us (A. G. M.). The Patellogastropoda are thought to represent the earliest surviving gastropod offshoot after the origination of torsion and to harbour many plesiomorphies for the group (reviewed in Haszprunar 1988a; Ponder & Lindberg 1996). It was thus exciting when possibly the most primitive living lineage of patellogastropod was discovered at hydrothermal vents and hydrocarbon seeps in the form of the Neolepetopsidae (McLean 1990a). Based primarily upon the existence of articulating radular teeth within the Neolepetopsidae, previously only known from Lepetopsis (Mississipppian to Triassic), McLean (1990a) erected the new suborder Lepetopsina (Neolepetopsidae + Lepetopsis) to represent the most basal split, of Palaeozoic origins, within the Patellogastropoda. Fretter's (1990) anatomical investigations favoured close relation of the Neolepetopsidae to the Acmaeidae (fossil record beginning in the Triassic; Tracey et al. 1993). Molecular investigations using 16S rRNA sequences (R. Guralnick pers. comm.) and 18S rRNA sequences (A. G. M. with M. G. Harasewych, in prep.) support Fretter's (1990) contention of acmaeid affinities and thus reject Palaeozoic origins of the Neolepetopsidae, although possible Mesozoic origins cannot be addressed without additional taxonomic sampling of the Acmaeidae and close relatives.

279

Based upon novel shells and radula, McLean 0988) designated the new superfamily Lepetodriloidea for two new families (Gorgoleptidae and Lepetodrilidae) of the Vetigastropoda, represented by Lepetodrilus and Gorgoleptis. Lepetodrilus is now represented by eleven species and the genus is known from hydrothermal vents of the eastern and western Pacific and the mid-Atlantic Ridge. The superfamily is thus representative of a large, in situ radiation. The Vetigastropoda is equivalent to the traditional 'Archaeogastropoda' (Hickman 1988; Haszprunar 1993) and has a fossil record back to the Late Cambrian (Tracey et al. 1993). The internal anatomy of the Lepetodriloidea exhibits a primitive level of organization (Fretter 1988) but there are no known fossils. McLean (1988) provisionally ranked the Symmetrocapulidae (Jurassic to Early Cretaceous) as a sister-taxon and hypothesized entry of the Lepetodriloidea into hydrothermal vents during the Palaeozoic or Mesozoic. Phylogenetic reconstructions (e.g. Haszprunar 1988a) and molecular systematic investigations using 28S rRNA (Fig. 2) and 18S rRNA (A. G. M. with M. G. Harasewych, in prep.) support placement of the Lepetodriloidea within the Vetigastropoda, but are uncertain as to its exact position or phylogenetic novelty. One additional family (Clypeosectidae, placed in Fissurelloidea by McLean 1989a) and two subfamilies of the Scissurellidae (Sutilizoninae and Temnocinclinae, McLean 1989a) have also been described for the Vetigastropoda. Both the Clypeosectidae and Temnocinclinae lack collabral ridges, a presumed plesiomorphy unknown in fossil and Recent a r c h a e o g a s t r o p o d s (McLean 1989a). Both the Fissurelloidea and Scissurelloidea have Mesozoic origins, but the age of the vent endemic groups remains unclear. Molecular sequences of 18S rRNA support the scissurellid origins of Temnocinclis (Temnocinclinae), but cannot make any comment upon its phylogenetic novelty (A. G. M. with M. G. Harasewych, in prep.). The systematics of the Fissurelloidea-Scissurelloidea complex has a convoluted history and requires detailed investigation before the origins of vent endemics become clear. The Cocculiniformia (Cocculinoidea + Lepetelloidea) represents a diverse group of sulphidetolerant deep-sea families known from hydrothermal vents, hydrocarbon seeps, sunken wood, algae, and seagrass, as well as various invertebrate and vertebrate remains (Haszprunar 1988b; Marshall 1996). While one family, Pyropeltidae McLean & Haszprunar 1987, is endemic to hydrothermal vents, hydrocarbon seeps, and sunken whale bone, the overall

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systematics of the Cocculiniformia is unclear due to the relatively recent investigation of the group and the description of many new families. Species of the Cocculinidae and Pseudococculinidae are also known from hydrothermal vents. Both of these exist on biogenic substrata at bathyal and abyssal depths (McLean & Harasewych 1995)-1ike many deep-sea cocculiniforms, definition of endemicity relative to sulphide tolerance is unclear. Thus, the Pyropeltidae has its origins from a deep-sea endemic group that is adapted to exotic and mostly sulphide-influenced habitats, although it is unique in not requiring a substratum of biological origin (McLean & Haszprunar 1987). The few cocculiniform fossils are Cenozoic (mid-Paleogene to Recent; Tracey et al. 1993) and the phylogenetic origins of the group are unclear (e.g. Haszprunar 1988a; Ponder & Lindberg 1996). Much additional research into this extraordinary group of gastropods is needed before the origins of the Pyropeltidae can be understood. The family Provannidae, with published reports from hydrothermal vents and hydrocarbon seeps, also represents a large, in situ radiation. Four genera and nineteen species are known. In addition, J. Voight and A. War6n (pers. comm.) report the first discovery of Provanna in association with sunken wood. Upon close examination of Provanna, War~n & Ponder (1991) placed it in the new family Provannidae with close affinities to the deepsea endemic Abyssochrysidae, represented only by Abyssochrysos. Both families exhibit plesiomorphic neotaenioglossan features and probably r e p r e s e n t the s t e m - g r o u p of the Caenogastropoda (War~n & Ponder 1991; Tracey et al. 1993). War~n & Ponder (1991) place the Provannidae with Abyssochrysidae in the Loxonematoidea (Ordovician to Jurassic, Tracey et al. 1993), after Houbrick (1979). The deep-sea thus harbours two relic families of the otherwise extinct, shallow-water Palaeozoic superfamily L o x o n e m a t o i d e a - P r o v a n n i d a e (vents, seeps, sunken wood) and Abyssochrysidae (normal deep-sea). Specimens of both Provanna and Abyssoch~3,sos are additionally known from an Eocene hydrocarbon seep community (Squires 1995; Goedert & Kaler t996). It should be noted that four species of the Provannidae have been dredged presumably away from hydrothermal vents and hydrocarbon seeps, but that the exact nature of their habitat remains unknown (War~n & Bouchet 1986; 1993). One of the most exciting gastropods discovered at hydrothermal vents was Neomphalus fretterae, generally considered a living fossil of

an otherwise extinct major radiation of the Palaeozoic-Mesozoic (McLean 1981; Batten 1984a). This gastropod presents characteristics plesiomorphic for the Gastropoda predominant in the archaeogastropod grade of organization but uniquely combined with characteristics previously considered synapomorphic for higher gastropods (McLean 1981; Fretter et al. 1981). Neomphalus belongs to an endemic hydrothermal vent suborder, informally known as the Neomphalina, that now consists of two superfamilies, three families, 17 genera, and 33 species. Species are known from hydrothermal vents of both sides of the Pacific and new species are currently being described from the midAtlantic Ridge (A. War+n, pers. comm.). The Neomphalina thus represents one of the largest in situ radiations at hydrothermal vents. McLean (1981) proposed affinity of the group to the extinct, shallow-water Euomphaloidea (Ordovician to Permian, Tracey et al. 1993) and proposed the new suborder Euomphalina to include these fossils and the newly-discovered vent representatives. While this association has been contested (Batten 1984a,b), phylogenetic novelty within the Gastropoda equivalent to Palaeozoic origins has generally been supported (e.g. Haszprunar 1988a; Ponder & Lindberg 1996). However, opinions on the group vary from phylogenetically novel mono-, poly-, or paraphyletic archaeogastropods (McLean 1981; 1989b, 1990b; Haszprunar 1988a; Ponder & Lindberg 1996; Salvini-Plawen & Steiner 1996), recently derived products of rapid morphological change (Hickman 1984, but see Haszprunar 1989), or an ancient group within the mesogastropod grade (Batten 1984a,b). Its anatomical novelty and the complete lack of fossils with similar shells made the Neomphalina a perfect candidate for molecular systematic investigations by one of us (A. G. M.). Full details will be presented elsewhere, but below we summarize the methods and results. Five previous studies had used partial 28S ribosomal RNA sequences from two variable domains to study the phylogeny of the Gastropoda (Emberton et al. 1990; Tillier et al. 1992, 1994, 1996; Rosenberg et al. 1994). Sampling of the early archaeogastropod lineages, likely relatives of the vent Neomphalina, was incomplete in these studies and each illustrated that use of a single, short domain provided limited resolution of Palaeozoic radiations. Fifty-three new DNA sequences of the D1 and D6 domains of the 28S rRNA gene were obtained from a representative sampling of the Gastropoda (plus outgroups) by polymerase chain reaction amplification followed by automated D N A sequen-

ANTIQUITY IN THE MODERN VENT FAUNA cing. These were combined with some of the previously published sequences to result in r D N A / r R N A sequences of both domains for three outgroups (two Polyplacophora, one Bivalvia) and 32 ingroup taxa (three Patellogastropoda, five Neritimorpha, five Neomphalina, eight Vetigastropoda, five Caenogastropoda, six Heterobranchia). Sequences of Cephalopoda (the nearest outgroup) and Cocculiniformia (possibly important ingroup) could not be obtained. Phylogenetic trees were reconstructed using parsimony, distance methods, and maximum likelihood. Full details of the analyses and results are presented in Fig. 2. In brief, phylogenetic reconstructions strongly supported monophyly of the sampled gastropod (sub)orders: Heterobranchia, Caenogastropodm Neritimorpha, Vetigastropoda, and Neomphalina. Although the sampled Neomphalina lacked one family (Neomphalidae), the two superfamilies were monophyletic. Additionally, Melanodrymia was not paraphyletic to the remaining Neomphalina, as suggested by Salvini-Plawen & Steiner (1996) because of lack of the synapomorphy of skeletal rods in the ctenidia (gills). In terms of evolutionary implications, the four non-vent groups have origins ranging from the Late Cambrian to the Devonian and the sampled families from the Early Triassic to mid-Jurassic. The Neomphalina does not belong within these early-mid Mesozoic radiations of the other major gastropod lineages and is thus statistically supported (i.e. bootstrap values) as an early-mid Mesozoic relic. Investigations of 18S rDNA further indicate similar novelty of the Neomphalina from the Patellogastropoda and Cocculiniformia (A. G. M. & M. G. Harasewych, in prep.). Its actual origins are probably Palaeozoic as evidenced by the small internal branch lengths of Fig. 2. Ribosomal D N A investigations may never be able to resolve this portion of the phylogeny as the majority of lineages needed to split internal branches are extinct - the majority of extant marine prosobranch gastropod families are of mid to lateCretaceous origins. Bivalvia. The bivalve taxa at vents are not diverse: only five families, three of which have only one or two species recorded. Sulphideoxidizing symbioses between microbes and bivalve molluscs have appeared in four lineages: Mytilidae, Vesicomyidae, Solemyioda and Lucinacea. The first two groups have flourished and diversified extensively at vents but, curiously, the lucinaceans have not yet been found. Despite the radiation of lucinaceans in the upper Palaeozoic (Reid & Brand 1986), the preadaptation of

281

symbiosis has not suited these animals to the vent environment. Lucinaceans are known at seeps (e.g. MacDonald et al. 1990), so the lack of sediments in most vent habitats may be a key deterrent as may be their apparent preference for very low sulphide levels (Fisher 1990). The only endemic bivalve group at hydrothermal vents and hydrocarbon seeps is the Bathymodiolinae, a subfamily of mytilid mussels represented by Bathymodiolus. While possibly having primitive gut anatomy for the Mytilidae (Kenk & Wilson 1985), nothing is known about its phylogenetic position or age within the family. The family itself is an old one: the Mytilidae radiated in the Carboniferous. The adaptation of byssal attachment allowed these filibranchs to invade hard substratum habitats. By examining allozyme and morphology variation in vent and seep Bathymodiolus, Craddock et al. (1995) did not find shared species between hydrothermal vents and hydrocarbon seeps, suggesting that faunal interchange may be limited to singular invasion events instead of continuous interchange. Genetic distances among Bathyrnodioh~s species indicated distant species deserving of multiple supraspecific taxa. Their phylogenetic analyses illustrated a pattern of invasion from shallow-water to deeper-water, with invasion of cold-water hydrocarbon seeps preceding invasion of hydrothermal vents. Thus, the Bathymodiolinae has probably had a long historical relationship with hydrothermal vents and hydrocarbon seeps, but how long remains unknown. The systematic relationships within the vesicomyids are in disarray at present (Boss & Turner 1980; Vrijenhoek et al. 1994). As currently used, the family comprises species of Vesicomya and Calyptogena. The vesicomyid radiation at seeps and vents is likely to be a Cenozoic phenomenon (Keen 1969; Morton 1996). Goedert & Squires (1993) present an unequivocal record from an Oligocene seep deposit. However, Calyptogena has also been noted from Cretaceous seep outcrops in Japan (Kanie et al. 1993 in Naganuma et al. 1995) where early Cenozoic sites are also known. The acquisition of symbionts and the adaptation of the foot for sulphide uptake appear to be major contributors to the radiation of this group (Reid 1990). Many more species are known from vents and seeps than elsewhere; possibly the vesicomyids originated in these habitats. While the Solemyidae is not well represented at vents, they belong to a group that has a record to the early-mid Palaeozoic. No adaptive barrier to their early entry into vents seems to exist, yet this group is predominantly found in other

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sulphide-rich environments (Reid 1980; Felbeck 1983). Of the 22 bivalve species (Tunnicliffe, McArthur & McHugh, in prep.), only four contain no symbionts: two peripheral mussels, a nuculinid at a sedimented site, and a pectinid. The pectinid, Bathypecten, was described as primitive when discovered ( S c h e i n - F a t t o n 1985). A second species discovered outside the vent environment was placed in the same new genus (with some uncertainty) and an analysis of all the pectinid genera placed Bathypecten as the least derived (Schein-Fatton 1988). The vent species retains the most plesiomorphic feature calcitic prisms on both v a l v e s - t h a t is otherwise known only in the Pectinacea before the Permian. Schein-Fatton (1988) calls for more work to place the second species and discusses the 'relict' character of the vent species. Perhaps this genus is an example of a relatively recent move of a refuge taxon away from the vent habitat.

Hemichordata Although most species of enteropneusts are known from intertidal or shallow water areas, the monospecific family Saxipendiidae has been reported at hydrothermal vents of the Galapagos Rift (Woodwick & Sensenbaugh 1985). All published reports of Saxipendium coronatum are at hydrothermal vents, but Van Dover & Hessler (1990) indicate it is a non-vent species occurring in greater abundance in the periphery of diffuseflow vent fields than elsewhere in the deep-sea. We are additionally aware of unpublished reports of it away from vents. As stated by Woodwick & Sensenbaugh (1985), little consideration has been given to the systematics of the Enteropneusta since the beginning of the 20th century. Although the Saxipendiidae is probably not a vent endemic, we include it in our list until its distribution is better documented.

vent, neither monoplacophorans nor brachiopods are part of modern vent communities. This suggests they faced competitive exclusion at vents from invading bivalves and archaeogastropods as they did in other marine communities during the Palaeozoic (Sepkoski & Miller 1985; Campbell & Bottjer 1995). Many modern vent inhabitants have congenerics outside vents, indicative of recent invasion. For example, the common vent squat lobster is Munidopsis spp., a genus well represented in deep-sea environments, while the vent hydroid Candelabrum shares its genus with shallow-water relatives (Segonzac & Vervoort 1995). Some peripheral vent species have succeeded in tolerating mild sulphide conditions in relatively recent penetration of the habitat (i.e. Beck 1996). Others may, with closer inspection, prove not to be endemic at all as more effort is spent on exploring other deep-sea environments. Recent discoveries away from sulphide-rich habitats of ancient groups originally only known from vents, and vice versa, are particularly interesting. Examples include the least derived pectinid discussed above, the discovery outside vents of erebonasterid copepods, and the discovery of loxonematoid gastropods in vents, seeps, sunken wood, and the 'normal' deep-sea. Have these primitive animals been secluded in the deep-sea and recently moved into vents or have the vents provided a refuge from which they recently ventured forth? Have they possibly existed in both habitats for a long time? Further exploration in the 'normal' deep-sea will, hopefully, provide more specimens for systematic investigation of the origins of these primitive animals. The vent habitat is clearly not impenetrable, but we have presented evidence for the refugial survival of some ancient lineages. Next we present our views on possible mechanisms responsible for their survival. We present these as hypotheses testable by future systematic and biogeographic investigations.

Faunal origins Like many marine communities, the assemblage of animals at hydrothermal vents has mosaic origins and has probably experienced a continuum of invasion throughout the Phanerozoic - the groups mentioned above include those of Palaeozoic, Mesozoic, and Cenozoic origins. Fossil vent communities are very rare, but Little et al. (1997) document the association of m o n o p l a c o p h o r a n molluscs and inarticulate brachiopods with a Silurian vent assemblage. Although A. War~n (pers. comm.) reports a modern monoplacophoran near a hydrothermal

Physiological barriers Evidence for the importance of barriers to invasion for refugial survival of lineages comes from two sources: (i) hydrocarbon seeps, while similar, present less severe barriers to invasion and the seep fauna is much less endemic, although some interesting endemics do occur (e.g. the vestimentiferan family Escarpiidae, Jones 1985); (ii) most vent species are endemic yet ~50% of

ANTIQUITY IN THE MODERN VENT FAUNA genera are found outside of sulphide-rich environments, suggesting that invasion must be accompanied by (taxonomically identifiable) adaptation to the physiological extremes. These barriers presumably excluded some groups from invading, providing a refuge for others. Three features of the habitat are relevant: its isolated, island nature, the presence of sulphide, and the high variability of parameters such as dissolved gases and temperature. Sulphide is toxic to aerobic organisms because of its inhibitory actions on cytochrome oxidase, a vital enzyme in the aerobic respiration process. Encounters with hydrogen sulphide are not uncommon in the marine environment and many mechanisms to cope with it are described (e.g. Somero et al. 1989; Vismann 1991). Behavioural and physiological adaptations include avoidance, importation of oxygen, detoxification, immobilization of sulphide and facultative anaerobiosis. Plenty of time has been available for organisms in sulphide-rich sediments to develop pre-adaptations for the sulphide-rich waters of hydrothermal vents. For example, species of N e r e i s polychaetes are found in both sulphide-rich sediments (Miron & Kristensen 1993) and at vents. The levels of sulphide encountered by vent animals are probably higher than most sediment dwellers: field determinations around animals are up to 300#M (Johnson et al. 1988) but lab experiments show much higher tolerances (Childress & Fisher 1992). As sulphide readily diffuses across epithelial surfaces, the major adaptation of vent animals to sulphide tolerance seems to be oxidation of the compound to a "safer' form (Childress & Fisher 19921). In addition, most aerobic metazoans in sulphide-rich habitats tend to have a physical barrier over most of the body: mucus-rich cuticle, shell or carapace. As it is unusual to find more "exposed' animals such as echinoderms, coelenterates, or bryozoans, sulphide may present a physiological barrier to vent entry that excludes some common deep-sea groups. The isolated, island nature of hydrothermal vents may not present a barrier to invasion as phenomenal dispersal capabilities are wellknown for marine larvae (Scheltema 1986), even for those animals without a larval stage (Johannesson 1988). Gene flow studies of hydrothermal vent organisms support this view (Vrijenhoek 1997). The complication of an ephemeral habitat appears more challenging but many animals can also find deep-sea windfalls such as wood or large carcasses (Gooday & Turley 1990). Simi-

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larly, many vent animals appear to colonize nascent vent sites quickly (Vrijenhoek 1997). An unusual feature of vents is the high rate of change of many environmental variables both spatially and temporally (Tunnicliffe 1991; Childress & Fisher 1992). A wide tolerance range is essential and may represent one of the unique adaptive requirements at vents (Gorodezky & Childress 1994). This feature, combined with high sulphide levels and the disparate nature of the habitat, defines a set of habitat conditions that forms a formidable barrier to entry by poorly adapted animals attracted by the high productivity. Tunnicliffe et al. (1996) discuss the concept of an ancient sulphophilic fauna that includes the common element of vents, seeps, and deep-sea organic remains. While the majority of endemic groups are restricted to vents, several are found in other sulphide-rich habitats. Evolution of sulphide tolerance in one habitat may thus be a threshold adaptation for invasion of related habitats. For example, vents and seeps differ in the delivery mechanism and flux of sulphide; water column sulphide is very low at seeps (Scott & Fisher 1995). Gradients and physio-chemical conditions are less extreme at seeps resulting in a less clear boundary with surrounding faunas; entry appears less restricted. The net result appears to be much greater encroachment upon seep sites by 'normal' deep-sea animals. Seeps on accretionary sediments are probably sites of long-term emissions compared to vents (Aharon et al. 1997) and thus may encourage adaptation in the local populations. One might imagine adaptation to sulphide-rich seeps that subsequently provides adaptations to the more difficult vent habitat. Molecular studies indicate evolutionary linkages between vents and seeps (Williams et al. 1993; Vrijenhoek et al. 1994; Craddock et al. 1995) and further work along these lines will test whether an important part of the vent fauna holds a common ancestry with other sulphide-rich habitats. Several recently derived groups, such as the neolepetopsid limpets, nautiliniellid polychaetes, the many cocculiniform gastropod families, alvinocaridid shrimps, and some subfamilies of polynoid polychaetes have invaded multiple sulphide-rich habitats and may provide excellent opportunities to examine the importance of sulphide tolerance and other adaptations to the successful invasion of hydrothermal vents. Both vent and seep animals go beyond sulphide tolerance to exploitation of symbioses with sulphide-oxidizing microbes. The resulting biomass far exceeds any other sulphide-rich habitat. Symbiotic relationships form a core of

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the trophic structure at vent and seep sites and symbiosis probably represents another threshold adaptation in many groups, possibly acquired first in other sulphide-rich habitats, for invasion of hydrothermal vents and hydrocarbon seeps. Evasion of mass extinction

Mass extinctions play an enormous role in the changing profiles of evolutionary faunas. During very short geologic intervals, drastic biotic shifts appear. Mass extinction in the deep-sea would argue against the existence of refugial ancient groups. However, some groups found in vents and seeps do appear to represent ancient lineages that have survived independently in these sulphide-rich refugia. These groups include the basibranch pogonophorans, bathymodiolinine mussels, dirivultid copepods, and provanhid gastropods. Perhaps the very nature of sulphide-rich environments has provided some measure of protection from certain causes of extinction. Other ancient groups appear to require the severity of the vent environment for protection, possibly as a result of the greater openness at other sulphide-rich environments. These groups include the diverse groups of barnacles, many basibranch pogonophoran families, axonobranch pogonophorans, neomphalinid and l e p e t o d r i l o i d e a n g a s t r o p o d s , ecbathyriontid copepods, archinomid and alvinellid polychaetes. Tunnicliffe (1992) suggested that extinction events had less impact in the vent environment than in most marine habitats. While there are many potential causal agents for mass extinctions, recent evidence concerning the terminal Permian and Cretaceous extinctions appears to converge on two proximal causes: a major change in ocean chemistry and a drastic reduction in primary productivity, respectively. The late Permian was a time of great carbonate accumulation (Grotzinger & Knoll 1995) that suggests unusual processes involving carbon dioxide. Knoll et al. (1996) build a scenario of ocean stratification and overturn that resulted in the largest metazoan extinction of the Phanerozoic. They emphasize the role of hypercapnia - metazoan poisoning by increased CO2 - possibly complicated by H2S presence as deep anoxic waters moved upward. However, any major increase in partial pressure of CO2 in seawater is likely to benefit the symbiont hosts at vents as they may be limited by available, diffusable carbon dioxide. The blood pCO2 is highly elevated in Riftia (Childress et al. 1993) and Calyptogena (Childress et al. 1991); R([tia appears to store CO2 (over 30 mM) in the blood

to exist 'in a state of compensated hypercapnic acidosis' (p. 1455, Toulmond et al. 1994). While these animals may represent the extreme of adaptation, vent animals appear to have adapted to high pCO2 levels in vent waters which, around tubeworms, is recorded at two orders of magnitude higher than the surrounding deep-sea water (Childress et al. 1993). CO2 concentrations can be so high that it bubbles out of solution at depth (Butterfield et al. 1990; Sakai et al. 1990). Adaptation to high pCO2 levels at hydrothermal vents may have allowed some animal lineages to survive through the endPermian and similar mass extinction events to the present day while going extinct elsewhere. The late Cretaceous extinctions resulting from bolide impact caused major climate changes (Hildebrand et al. 1991); decreased production probably ensued (Paul & Mitchell 1994; Levinton 1996). There is some evidence that surface primary producers were more strongly affected than deep water plankton (Milne & McKay 1982). Communities away from the continental shelf with their own in situ source of organic carbon through chemosynthesis may be expected to experience even lower influences of the impact 'fallout'. It should be noted that MacLeod et al. (1997) present several challenges to the causative link between bolide impact and the late Cretaceous extinctions- biotic change may have been going on prior to the bolide impact. While many extinctions during the late Cretaceous may have been products of prolonged environmental change, MacLeod et al. (1997) did find evidence of a rapid extinction of marine plankton at the K-T boundary. We suggest the major reasons that mass extinctions had less impact in the hydrothermal vent environment are firstly, that physicochemical state shifts in the ocean are already accommodated within the highly heterogeneous vent habitat and, secondly, that the vent inhabitants have a much lower dependence upon solar-based primary productivity. As radiation of new groups relates very strongly to the eradication of incumbents, displacement of long-time inhabitants which have well-developed adaptations to hydrothermal conditions would be much harder at vents if the inhabitants exhibit immunity from mass extinctions. However, this may not apply to the entire habitat - it is clear that some vent niches are highly variable and susceptible to a variety of influences such as major shifts in oxygen availability or suspended particulates from the surface. A major extinction agent from which vent communities would not be immune is anoxia. Invertebrate metabolism depends on oxygen no

ANTIQUITY IN THE MODERN VENT FAUNA matter how well adapted the animals are to fluctuating concentrations and hydrogen sulphide incursions. Deep-sea anoxic vents have been hypothesized to result from high-latitude warming and changes, including reversals, in oceanic circulation (Wignall & Twitchett 1996: Speijer et al. 1997). The support for deep-sea anoxia is u n c l e a r - models vary among highly reduced oxygen (e.g. Kennett & Stott 1991), global anoxia or dysoxia (e.g. Bralower et al. 1994; Wignall & Twitchett 1996; Isozaki 1997; Speijer et al. 1997), and regional anoxia (e.g. Galil & Goren 1994). Hydrothermal vent communities are found in the upper third of the ocean (mostly above 2200 m) but it is unclear if historical anoxic events have risen out of the abyssal depths. Global anoxic events that included the deep-sea have been proposed for the P e r m o - T r i a s s i c b o u n d a r y (Wignall & Twitchett 1996; Isozaki 1997), early Toarcian (i.e. Jurassic; Little 1996; Aberhan & Ftirsich 1997; Hori 1997), and late Palaeocene (Kennett & Stott 1991; Speijer et al. 1997). The impact at bathyal depths of other global anoxic events (e.g. Harries 1993) is less clear. Vent communities were unlikely to have been immune to deep-sea anoxic crises and the persistence of Palaeozoic and Mesozoic lineages at modern vents may require additional explanation.

Conclusions

The hydrothermal vent fauna has mosaic origins, with a component representative of ancient refugial survivors of the Mesozoic and Palaeozoic. The evolution of sulphide tolerance and chemoautotrophic symbioses may have led to protection from some causes of extinction. Other sulphide-rich habitats, such as hydrocarbon seeps, thus may have additionally acted as deep-sea refugia. Vents are one of the best explored deep-sea ecosystems and the fauna one of the most understood. Indeed, our examination of its origins was hampered more by a lack of understanding of other communities than of the vents themselves. We suspect that the interaction of sulphide with marine communities has been an important evolutionary dynamic throughout the Phanerozoic and we greatly anticipate exploration of additional deep-sea habitats. Inclusion of taxa from other habitats is needed for both phylogenetic and biogeographic comparisons. In addition, many vent groups have not been examined since their original description. Future applications of phylogenetic frameworks to the examination of deep-sea biogeographic patterns, combined with

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new discoveries, will provide the means to test our ideas about the evolutionary origins of the hydrothermal vent fauna. Contributions to our database of hydrothermal vent associated molluscs were made by A. Lesicki. Numerous systematists have contributed information and we have benefited from communications with W. A. Newman, A. War6n, and K. Fauchald. Insights into mass extinction and anoxia were received from C. Little and this manuscript was much improved by comments from two anonymous referees. Assistance from the Smithsonian Institution Libraries, A. Schulze, and L. Franklin is gratefully acknowledged. NSERC Canada and the University of Victoria provided research funding to V. Tunnicliffe. A. G. McArthur was supported by a Smithsonian Institution Postdoctoral Fellowship and the National Museum of Natural History's Laboratory of Molecular Systematics.

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Index

Note: Page numbers in italic type refer to illustrations, those in bold type refer to tables.

Abyssochrysos 280 accommodation zone 56 accretionary prisms 220, 235 actinolite 89, 90, 135, 146, 147 Adria 220 advection 99, 139 Aegir Rift 77 Afro-Arabia 63 airgun arrays 21 A1/Ti ratios 208, 210 albitization 58, 93, 160, 195 alteration low temperature 115, 159, 171, 172 polyphase 145, 172 alteration front 132 alteration haloes 115 alteration modes 108 alteration petrography 108 alteration pipes 155 alteration stages Juan de Fuca Ridge 114, 115, 116, 118 Troodos ophiolite 142 alteration zones oceanic crust 82 Pitharokhoma pipe 160 Troodos 129 Alvinella pompejana 260, 264, 265, 266 Alvinellidae 263, 275 Alvinocaris 277 A M A R programme 1, 14 amphibole, vein 90, 94 amphibolite facies 83 amygdales 141 analcite 87, 88 Anarhynchia gabbi 260 Andaman Sea 62 andesites 160 anhydrite Bent Hill 186, 187 in greenschist facies 92 ODP Hole 504B 92 replacement of 93 spires 251 Sr in 90, 94 TAG mound 203, 211 anisotropy, gabbro 75 Annelida 275 anoxia 272, 284-285 Antalya Complex 233, 237 antigorite 75 antiquity hypothesis 272, 274, 275, 277 Arabian platform 53 aragonite 86

arc basins 219 archaeogastropods 263 Archinome 276 Argolis Peninsula 233, 236 Aroania 222, 223, 224 As 195, 196 As/Fe ratios 203 asthenosphere, upwelling 63 atacamite 203, 205, 209, 210, 213 Atlantis II Deep 212 Atlantis II Fracture Zone 31 bathymetry 73 seismic velocity model 74 transverse ridge 76 axial calderas 17 axial depth 17 axial magma chamber 17 models 22 axial valley Gorda Ridge 189 Mid-Atlantic Ridge 29 segment OH1 5 segment OH3 8 Azema 260 Azores Triple Junction 1 back-arc basins 219 bacterial cells 255, 264 Baer-Bassit units 233 barite 191,242, 245, 246, 248, 252 Barlo 260, 265 barnacles 272, 277, 278 phylogeny 277 basal reflector, absence of 18 basalt alteration 88, 209, 212 clastic 209 and fluid fluxes 212 glassy 10 massive flows 106, 184 MgO content 17, 3 I, 58 picritic 25 within-plate 233 see also MORB basaltic glass 141 basement, burial of 120, 122 basement fluids, Sr ratios 85 basement temperatures, Middle valley 178 basement-sediment interface, temperatures 103, 106, 122 bathymetric maps, M A R segments 3 bathymetric sections, segments OH1 and OH3 7 Bathymodiolus 267, 281

294

INDEX

Bathypecten 282

Bayda 260 Benioff zone 62 Bent Hill massive sulphide 177, 178, 188 lithostratigraphy 179, 180 rock types 182-183 Besshi-type deposits 186 Bi 195, 196, 257 tellurides 247 bivalves 262, 263 Bivalvia, endemicity 274, 281-282 black smokers Archaean 241 EPR 20 fluid analysis 203 fluxes 129 and ochres 218 REE composition 153, 154~ 169 Sr ratios 86, 87, 93, 94, 136, 148 temperatures 90 Blanco Fracture Zone 189 block faulting 100, 106 block rotation 54 boudinage 46, 54 Bouguer anomalies, see mantle Bouguer anomalies boundary layer, temperatures 148 brachiopods 243, 260, 263, 264, 267 breccias hyaloclastite 86 ODP Hole 896A 82 ODP Leg 168, 106, 107 Sr ratios 85 striated 6 sulphide 180, 242 brittle structures 54 Brunhes Epoch 178 Buribuy 262 Buried Basement transect 100, 103-106 Caenogastropoda 280 calcite, secondary mineral 86, 209 calcite compensation depth 265 calderas, axial 17 Calyptogena 263, 281,284 Candelabrum 282 carbon dioxide, and mass extinction 284 carbonate accumulation 284 authigenic 190 biogenic 203 dissolution rates 265 secondary minerals 88, 110, 116, 121 Carrick Luz shear zone 33-39 principal features 35 Cascadian Basin 103 cataclasite 36 celadonite in alteration assemblages 85, 110 formation temperature 120 in oxidation haloes 115, 116 Troodos 142 Central Hill 189 lithostratigraphy 189, 190

location map 190 rock types 192 sulphide deposits 191 chalcophile elements 203, 208, 210, 213 chalcopyrite Bent Hill 181, 185 conduit lining 251,257 disease 247 in umbers 52 Urals 242, 245, 246, 247 characteristics, MAR segments 4 chert breccias 223 cherts, Mn-rich, see Mn-chert chimney spires black smokers 241 fluid chemistry 255 fluid inclusions 243, 250, 257 fragments 246, 247, 251 isotope chemistry 249-240 mineral chemistry 248-249 mineralogy 245 248 sampling 243 zonation 245-248, 249, 255, 256, 257 chimney walls 251,254 chlorite 92, 93, 115, 119, 146 formation temperature 169 chlorite/smectite 110, 115, 119. 143, 160, 163-169, 171 Chorocaris 277 C1 concentration, pore fluids 193 Claire Seamount 8 clay minerals alteration products 209 Sr ratios 86 Cleft segment 31 clinopyroxene, recrystallization 89, 90 Clipperton 20 Co, enrichment 232 colloform banding 185, 227, 229 colonisation, vent sites 283 compaction length 25 complexation 153, 211 compositional layering 46 concentration-depth profiles 194 cones, pillowed 5 continent-ocean transition 77 copepods 276 copper, see Cu Costa Rica Rift black smokers 91 location map 82 Coverack 33, 39 cowlesite 165, 170 crack closure, depth relations 77 cross-cutting relations dyke-serpentinite 78 Kizildag 48 Troodos ophiolite 136, 141 crust, lower 44 Crustacea, endemicity 274, 276-279 crustaceans, Carboniferous 267 crustal age/distance from ridge 101 crustal ageing 155, 171 crustal denudation 53 crustal evolution 83

INDEX crustal temperature, and alteration 122 crustal thickness histograms 72 Lucky Strike segment 5 Na values 72 crusts, Mn-rich 217, 218 crystal content/depth relations 25, 26 crystals, gravity settling 26 Cu in massive sulphides 186, 188 supergene enrichment 209, 210, 236 Cu-Fe sulphides 181, 191 cunmlates Kizildag 46 Lizard ophiolite 33 in lower crust 17 Oman 25 Troodos ophiolite 154 decapods 276, 277 Deep Copper Zone 181 deformation, polyphase 45 Dellwood Seamount 219 detachment faults 56, 75 detachment surfaces 56 diabase Troodos 133 volume loss 129 diagenetic enrichment 219, 230 diapirization 56, 63 diffusive exchange 120, 139, 142 Dirivultus dentaneus 276 discontinuities, non-transform 2, 5 dolerite Juan de Fuca Ridge 106, 107 Kizildag 46 dolomite 185 down-core variation, metal ions 206 Drimos 222, 224 ductile flow 35 ductile shear 33, 35, 53 dunite 33, 46 Dy/Yb ratios, Kizildag 58 dyke injection and fissuring 17 Iceland 31 Kizildag 47, 49, 50 lateral 12 dyke/gabbro contacts, Kizildag 47, 49 dykes alteration minerals 143 diabasic 49, 83 gabbroic 35 multiple generations 49 Pitharokhoma 163, 170 vertical orientation 39 see also sheeted dykes earthquakes 52 East Pacific Rise bathymetry 18 cross-section 47 hydrothermal signature 91 magma chamber 52

295 rare-earth elements 212 seamounts 219 segments 19 seismic structure 26 study area 19 Eastern Manus Basin 219 Ecbathyrion prolixicauda 276 effluent temperature 127 ejecta deposits 6 electron microprobe analyses, see microprobe analyses elongation factor 275 emplacement modes, Juan de Fuca Ridge 106 Endeavour segment 100 endemism 272, 273, 274, 282 see also specific groups Eochionelasmus 277, 278 epidosite 129, 130, 133 epidote 90, 91, 93, 94, 147 eruption volumes 26 Escanaba Trough 177, 189, 267 Euomphalina, evolution 278, 280 Euphrosine rosacea 276 Explorer plate 177 extensional faults, continental 38 extensional tectonics, Kizildag 53-54 extrusive forms 12, 40 FAMOUS segment 1 geological interpretation 13 source 14 TOBI images 12 FARA-InterRidge programme 1 fault displacement, Carrick Luz 39 fault geometry 44 fault gouge 36, 136 fault orientation, Kizildag 53 fault scarps segment OH1 5, 6 segment OH3 10 fault surfaces, corrugated 31 fault zones, serpentinisation at 78 faults accommodation by 31, 40 depth of rupture 52 dyke-paralM 53, 54 fluid channelling 135 fluid down-welling 136 fluid up-welling 138 Kizildag 47, 51 as melt conduits 32, 39~ 40 Mid-Atlantic Ridge 44 mineralized 50 seismogenic 76 see also specific fault types faunal origins, vent communities 282 faunas deep ocean 271 vent-associated 241,243, 271 FAZAR cruise 2 feldspar alteration 90 ferromanganoan sediments 155 Figueroa 260 fissure swarms 5 fissures, density 17

296

INDEX

Flascarpia alv&ae 276 fluid channelling faults 135 flow and pillow margins 141 fluid flow buoyancy-driven 106 within-mound 203 fluid fluxes and basalt alteration 212 hydrothermal systems 90, 127 ridge flanks 99 ridge/flank transition 103 Troodos ophiolite 131, 145 variables 132 fluid inclusion studies 87, 243, 245, 250, 252 fluid-rock interactions 81, 99 kinetic controls 139 temperatures 83 fluids, see also hydrothermal fluids foliation planes 54. 75 fore-arc environment. Kizildag 62 formation temperatures, vein carbonates 88 fossil tubes 260, 265, 266 fossilization, at vent sites 264~266 fossils 259 identification problems 266-267 Sibay 262, 263 Yaman Kasy 263, 264, 265 fractional crystallization 61 Franciscan Complex 235 freezing horizon 25

gabbro-peridotite contact 36 gabbro-peridotite transition 71 gabbros Kizildag 46, 48, 54 leucocratic 46 plating 26 Troodos 154 Galapagos, Mn-rich deposits 218 galena 185 gangue minerals 191 Gastropoda, endemicity 274, 279 281 gastropods 260, 263, 267 gene flow 283 geochemical proxies 203 geochemical trends, Pitharokhoma pipe 160 geological sections, segments OH1 and OH3 6, 7 glasses, Troodos 163, 164, 167 goethite 191. 205, 209 Gondwana northwest margin 62 rifting 220, 271 Gorda Ridge spreading centre 189 Gorgoleptis 279 gossans 201,203, 204, 208, 212, 213 grabens, symmetric 29, 53 gravity lows, 'bull's eye" 5, 10 Greece, tectonic map 220 greenschist facies 38, 83, 92, 94, 128, 146 groundwater flow, fault control 31 growth faults 54 Guaymas Basin 195, 265

Gulf of Corinth 222 gyrolite 93 haematite, see hematite half-grabens 29 harzburgite 33, 46, 61, 154 Hayes fracture zone 1, 2 haystacks 8, 9 HEAT cruise 2 heat flow Buried Basement transect 106 hydrothermal systems 217 Hydrothermal Transition transect 103 ODP sites 504 and 896, 83 heat loss, ridge flanks 99 heat source, Troodos 147 heat transport, fluid flow 88 heavy rare-earth elements, Kizildag 58 hedenbergite 185 hematite 110, 115, 205, 246 Hemichordata, endemism 282 Hess Deep 44 HFS, see high-field strength elements high-field strength elements, Kizildag 58 highly altered rocks 94 hornblende, aluminous 38 horst and graben structures Kizildag 53 OHI segment 5 OH3 segment 10 hotspots, crustal thickness 77 HRE, see heavy rare-earth elements hyaloclastites Juan de Fuca Ridge 106, 107 OH1 segment 6, 8 hydrogen sulphide 283 hydrogenetic processes 230 hydrothermal alteration axial 153 and crustal age 122 evolution 119 late phase 12l low temperature 107 oceanic crust 75 open 120 restricted 120 sheeted dykes 88 hydrothermal chimneys collapsed 188, 191 OH1 segment 6 Urals 241 hydrothermal circulation 54, 115 diffuse 195 duration 148 geometries 129 high temperature phase 131 homogeneous 142, 148 Juan de Fuca Ridge 103 lateral 187 sediment control 196 Troodos 172 upper crustal layer 82 hydrothermal deposits, compositional variation 226

INDEX hydrothermal fluids composition 76, 118, 123 convection rates 122 downwelling 136 evolution 81,204, 213 lateral extent 187 low-temperature 217 off-axis 218, 233 pipe-like structures 147 rapid cooling 252, 253 rare earth elements in 204 temperature 82, 189 upwelling 129 see also fluid hydrothermal flux, estimates 128 hydrothermal reworking 186 hydrothermal systems structure 129 thermal structure 130 Hydrothermal Transition transect 100, 101-103 hydrothennal vents initiation by drilling 178, 187 intensities 17, 20 Lucky Strike segment 2 ODP deposit 178 hydrous component, Kizildag 63 hypercapnia 284 Iberian margin 77 Iceland, dyke injection 31 iddingsite 110, 115 illite 159, 169 imbrication 53, 221 immobile elements 160, 169 Impermeable horizons 187, 188 incompatible element ratios 58 interstitial fluids, geochemistry 187, 193 iron oxyhydroxides 110, 115, 120 isochron diagram, Troodos 144, 146 isocubanite 181, 191,247, 250, 251,257 isostatic uplift 53, 56 isotope chemistry, chimney spires 249-240 isotopic tracer transport model 128, 135 Izu-Bonin arc 229 jarosite 203, 213 jasper, Troodos 157, 170 Josephine ophiolite 32, 54, 78 Juan de Fuca ridge 31, 99 analytical techniques 107 cross-section 102 geological setting 100 lithostratigraphic logs 103 Middle Valley 177 Mn-rich deposits 218 juvenile arc model 62 K-Ar ages, alteration minerals 142 Kane fracture zone 31 Kane Transform 54 Kizildag Ophiolite 43 cross-sections 47, 55 evolution 62-65 geological map 45

297 location map 44 planar fabric elements 50 structural evolution 56 structure 46 K6mfirqukuru 52 Krasnogvardeyski 263 Kuroko deposit 241

La/Nd ratios 211 La/Sm ratios, Kizildag 58 Labrador Sea rift 77 Lamellibrachia 275 landslips 40 large-ion lithophile elements, Kizildag 58 laser ablation mass spectrometry 159-160 Latimeria 271 Lau Ridge 219 laumontite 91, 94, 143 lava lakes foundered 5 Lucky Strike segment 10, 14 Mid-Atlantic Ridge 2 lava tubes 9, 14 lavas age of 17 alteration 157 constructional 13, 14 Jurassic 222 lobate 5, 14 mixed-layer 163 rugged-flow 5 sheet-flow 5, 8, 13 layer 3 models 72 P-wave velocities 78 seismic velocities 3, 71, 75, 77 layering, cryptic 25 leached facies, Pitharokhoma 157, 160, 167, 169 lead, see Pb Lepetodrilus 279 Lepelopsis 279 lherzolite 33, 46 light rare-earth elements, Kizildag 58 LIL, see large-ion lithophile elements limonite 191 listric faults 50 lithological units, ODP Leg 168, 104, 105 lithosphere Neo-Tethyan 45 strength of 31, 39 lithosphere thickness, segment ends 40 lithostatic head 12 lithostratigraphic logs, Juan de Fuca ridge 103 Lizard Complex, map 32 Lizard ophiolite 32-33 block diagram 33 orientation diagrams 34 Ljeviha 263 location map 2 Loihi Seamount 219 low velocity zone 52 lowstands, Pleistocene 189 LRE, see light rare-earth elements Lucky Strike segment 1, 2, 5

298 geology 9 sidescan sonar results 10 source 13 TOBI images 8 LUSTRE 96 cruise 10 magma chambers East Pacific Rise 52 processes in 26, 44 magma lenses, axial 52 magma supply at fast-spreading ridges 52 at slow-spreading ridges 29 magmatic sources along-strike movement 10 Kizildag 58 single 14 magmatic-deuteric fluids 119 magnesium, see Mg magnetic anomalies ODP Leg 168, 101 seafloor spreading 77 magnetic measurements, resolution 75 magnetite 181, 185 Magnitogorsk 262 major element analysis, Kizildag 57, 59 malachite 52 Mamonia Complex 233, 237 manganese, see Mn mantle Bouguer anomalies, MAR 1, 5, 10 mantle decompression 74 mantle diapirs 10 mantle melting models, Kizildag 61 mantle/pluton contacts 54 marcasite 186, 191,247, 252 M A R F L U X programme 1 Mariana arc 229 Mariana Trough 61 M A R K area 3l, 44 core samples 54 cross-sections 47 M A R N O K area 31 mass extinction and carbon dioxide 284 evasion of 284-285 K-T 284 mass spectrometry, Pitharokhoma samples 159 mass wasting deposits 180, 202, 203 median ridge, segment OH1 5 median valleys 54 Melanodrvmia 281 melt body, AMC 22 melt layer thickness 25, 74 melt migration 32 melt topology 25 melt volumes 26, 61 melting models 58 Mendocino Fracture Zone 189 Menez Gwen segment 6 mesostasis 115, 116, 121, 142 metadolerite 49 metal distribution, in sulphide deposits 186 metal ions, down-core variation 206 metal remobilisation 203, 213, 230, 234

INDEX metamorphic rocks, Mn-rich 234 metamorphism, high-temperature 53 Metasequoia 271 metasolnatism 61, 63, 156, 157 Mg, in hydrothermal fluids 188 Mg-saponite 85 MgO content, basalt 17, 31, 58 microbial filaments 260 microprobe analyses, Juan de Fuca Ridge 107, 108, 109, 110, 111, 112, 116 Mid-Atlantic Ridge faults 44 TAG mound 202 variability along 1 Mid-Cayman Rise 53 mid-ocean ridges fast-spreading 17 slow-spreading 29 Middle Valley, Juan de Fuca Ridge 177, 178 mineral colour, and alteration 118 mineral precipitation 81 mineral recrystallization 128, 139, 143, 144 rates 150 temperatures 146 mineralization sequences 2 t 3 mineralized zones, Kizildag 52 Miura spinosa 276 Mn enrichment 206, 213, 218, 222, 235 fractionation 219 Kizildag 52 mining 222 precipitation 210, 217 Mn ores 233 Mn-cherts electron microprobe analysis 226-229, 228, 231 fractionation 234 geochemistry 223-226 hydrothermal origins 232 lamination 234 major and trace elements 225, 226, 227 Neo-Tethys 218, 220, 221 origins 229 petrography 223 photomicrograph 224 sedimentary log 224 SEM study 227, 229, 230, 232 settings 235 stratigraphic sections 223 ternary plots 229 Mn-nodules 234 Mo/Fe ratios 203 Modiomorpha mytiloides 263 Moho nature of 71 North Atlantic 76 ODP Hole 735B 75 petrological 33 thickness 52 monoplacophorans 243, 263, 265, 267 MORB 58 Greece 233 rare-earth elements 71, 154 Sr isotope ratios 84, 94

INDEX mordenite 165, 170 Mugodjarian zone 262 Munidopsis 282 mylonite gabbroic 33, 35 Kizildag 46 mylonite/gabbro contact 37 Na values, crustal thickness 72, 74 natrolite 87, 88 necking 54 Neo-Tethyan lithosphere 45, 62 Neo-Tethys basin 221-222 Mn-cherts 220 reconstruction 64, 233 seafloor spreading 236 subduction 234 see also Pindos Ocean Neobrachylepas 277 Neolepas 277, 278, 279 Neomphalus fretterae 280 Neopilina 271 Neoverruca 277, 278 Nereis 283 Newfoundland margin 77 normal faults 29, 39, 53, 178 Oasisia 263 obduction processes 45 oblique faults, Kizildag ophiolite 46 oblique-slip faults 55, 56 Ocean Drilling Program 33 see also ODP ocean stratification 284 OCEANAUT cruise 1 oceanic crust accretion of 17 alteration zones 82 comparison with Kizildag 52 formation of 44 models 72 thermal evolution 149 thickness 71 three-layer model 71 velocity structure 76 oceanic lithosphere, structural models 44 Oceanographer fracture zone 1 ochres and black smokers 217, 218 Cyprus 201 formation 202, 211 and massive sulphides 233 ODP Hole 504B 81 lithostratigraphy 84 location 83 sample 90 ODP Hole 735B, bathymetry 73 ODP Hole 896 82 location 83 ODP Leg 158 156 ODP Leg 168 99

299 location map 100 tabular data 104, 105 ODP Leg 169 177 location maps 178, 179 offsets, axial 52 OH I segment 1, 5 section 6, 7 OH3 segment 1 section 6, 7 Oktyabrsk 262 Oman, see Semail ophiolite belts Greece 233 peri-Arabian 46 ophiolites Apennine 234 fossils 260 harzburgite-type 77 Josephine 32, 54, 78 Kizildag 43 lherzolite-type 77 Ligurian 234 Lizard 32-33 Mn deposits 235 Oman, see Semail Pindos 201 Semail 25, 46, 201,205, 233, 260 as structural analogues 44, 71 Tethyan area 219 Troodos 46, 62, 63, 127, 201,260 Zambales 260 ore bodies Yaman Kasy 242, 251 zoning 210 orogenic belts 232 Othris 233, 236 oxidative haloes 113 oxyanions 203 oxygen isotope fractionation 75 oxygen isotope studies 87, 88

P-wave velocities, layer 3, 78 P/Fe ratios 203, 208, 209, 210, 213 Pabulum spathiforme 278 Pacific seamounts Fe-rich sediments 219 rare-earth elements 212 Pangaea, break-up 62 Papuan peninsula 63 paragenesis 141 Parah, mella sulfinicola 265 partial melting, seismic velocities 23 passive margins 233 Patras 221 Pb, in massive sulphides 186, 195, 196 peak-to-trough anomalies 5 Peloponnese 220 cross-section 222 geological map 221 peridotite feldspathic 46 partial melting 23 serpentinised 33, 44, 52, 56, 71, 75, 77

300

INDEX

Troodos 154 Peristerka 260 permeability boundary layer 129 and crustal structure 103 permeability-depth relations 77, 82, 94 petrogenesis, Kizildag magmas 58 phillipsite 115, 116 phyllosilicates 110 rare-earth elements in 211 Pico offset 1 pillow lavas FAMOUS segment 14 fluid flux in 130 Juan de Fuca Ridge 106 Kizildag 51, 52, 55 OHI segment 6, 8 OH3 segment 8 Sr isotope ratios 131 Pindos Ocean 221,222, 234, 236 Pindos ophiolite 201 Pitharokhoma alteration pipe major and trace elements 161 map 157 mineralogy 156 rare earth elements 162, 163, 164. 165, 166, 168 samples 158, 159 plagioclase, recrystallization 93, 195 plagiogranite 46. 154 plate separation and magma supply 53 role of faults in 29, 40 plesiomorphism 272, 275 plume deposits, umbers 218 plume fallout 202. 203. 204, 210 plutonic rocks, Kizildag 46 plutons, gabbroic 71 PO1 segment, see Lucky Strike segment Pocklington rise 63 Pogonophora, endemicity 274-275 Polychaeta. endemicity 274, 275-276 polychaetes 260, 263, 265, 267 pore fluids 177, 187, 193, 195, 196, 197 pore-filling, anhydrite 92 porosity boundary layer 129, 138 fault zones 188 porosity-depth relations 82 prehnite 93 primitive mantle 61 Provanna 280 pyrite alteration zones 116 analyses 254 Bent Hill 185, 191 bladed 252, 257 honeycombed 245, 256 massive 246 TAG site 203 Troodos 156, 167 in umbers 52 Urals 242, 245, 246 pyrite-marcasite 181, 185 pyrolusite 227

pyrrhotite 181, 184, 191 quartz, alteration mineral 110, i 16 quartz veins 90, 91 radiolarians 213, 222. 230, 235 rare-earth elements in alumino-silicate phase 210 black smokers 153, 154, 169 co-variation 211 East Pacific Rise 212 fractionation 204, 211,213 Kizildag 58 mobility 153, 170, 171 MORB 71 Pacific seamounts 212 partitioning 170 in phyllosilicates 211 Pitharokhoma 159, 163, 164, 165, 166, 168 ranges 212 in seawater 154, 167, 204 as tracers 204 Rb-Sr ages, alteration minerals 142 Rb-Sr evolution 138 reaction zone 83, 94 recharge cooling effect 147 pervasive 138 recharge fluids 85, 92, 93 geometry 128, 129 in hydrothermal flow 130 recharge flux 136 recharge zone 84 Red Sea 212 redox conditions 210, 2l 1 reflector, axial magma chamber 17 refugia 268, 282, 284 relic taxa 259, 271,275 replacement deposits 156 research vessels Atlantis H 204 Discovery 74 l" Atalante 2 Nadir 1 resistivity 188 retrograde solubility 92 rhodochrosite 235 rhyolite 242 ridge axis, cross-section 38 ridge crests 17, 44 ridge flanks fluid fluxes 99 heat loss 99 hydrothermal circulation 88 ridge segmentation 56 Ridgeia 260, 263, 265, 266. 275 rift valley segment OH1 5 segment OH3 5 Riftia pachyptila 260, 264, 266, 284 rifting 53 Gondwana 220 Mediterranean 62 rifts, extinct 77

INDEX Rimicaris 277 RNA, sequence analysis 275, 278, 279, 280-281 Rough Basement transect 100, 106

Safyanovka 262 sakalavites 52 Sakmara zone 242 salinity, fluid inclusions 250 sampling bias 84 Santalema miraseta 276 saponite alteration phase indicator 110, 119, 120 association with carbonate 88 association with chlorite I 15 association with pyrite 116 formation temperature 121 Sr signature 86, 142 see also Mg-saponite Saxipendium coronatum 282 scavenging mechanisms 170 Scillaelepas 278 seafloor spreading Kizildag ophiolite 46 magnetic anomalies 77 Neo-Tethys 236 seamount chains offset 31 segment OH1 5 seamount complex, Lucky Strike segment l0 seamounts fault-associated 31, 40 flat-topped 31 Mid-Atlantic Ridge 30 off-axis 219 seawater downward diffusion 211 entrainment 203 REE 154, 167, 204 Sr ratios 86 seawater penetration 77, 82 secondary minerals 81, 82 fluid composition 94 Juan de Fuca Ridge 110, 117 progressive change in 123 sheeted dykes 88, 89 Sr ratios 86, 87 textural variations 118 sediment hills 189 sedimentary cover Gorda Ridge 189 Juan de Fuca Ridge 103 Kizildag 53, 56 OH1 segment 5 sedimentation rates, Juan de Fuca Ridge 180 sediments alteration 190, 195, 196 interstitial 167 seep faunas 272, 274, 275, 276, 280, 282, 283 segment ends lithosphere thickness 40 lithospheric strength 31 segment evolution 14 segments East Pacific Rise 27

fourth-order 19, 27 lengths 29 segregation, fluid phase 193 seismic data, possible errors in 21 seismic profiles, East Pacific Rise 19, 20 seismic reflectors axial magmatic chamber 17-28 crustal 44 deep 39 seismic velocity layer 3, 71, 75, 77 melts 22, 23 Semail Ophiolite 25, 46, 201,205, 233, 260 serpentinisation 33, 56 conditions for 75-76 Josephine ophiolite 78 tectonic settings 76-77 serpentinisation front 71, 77 serpentinite, Troodos ophiolite 154 shear, sense indicators 54 shear strain 39 shear zones fluid channelling 135 gabbro-filled 32, 35 hydrated 37, 39 Kizildag 46, 49, 51 seismic reflectors 40 sheeted dykes hydrothermal alteration 88, 89, 131, 145 injection processes 26 Josephine ophiolite 54 Kizildag 47, 49, 50 Lizard 33 ODP 504 and 896, 83 Sr ratios 86, 93, 130, 139, 147 Troodos 54, 133, 155 Sibay, fossils 262, 263 SIGMA cruise 2 silica biogenic 232 precipitate 203 sill/sediment complex 181 sills axial magma chamber 17, 25, 26 basalt 192 Skouriotissa 209, 212, 213 slab window 63 slickensides 50, 54 slumping 202 smectite, Pitharokhoma 163, 164, 170 Snake Pit vent field 171 sonar, side-scan 1, 2, 10, 30 Southern Explorer Ridge 212, 213 Southwest Indian Ridge 29 melt migration 32 shear zones 54 tectonic stretching 44, 53 Sovanco transform fault 177 sphalerite Bent Hill 181, 185, 186, 191 collomorphic 247 dissolution 203 Urals 242, 245, 252 spreading centres 54, 63

301

302

INDEX

spreading rates Gorda Ridge 189 Juan de Fuca Ridge 100 Southwest lndian Ridge 74 Sr budget, global 94, 127 Sr exchange, timescale of 89 Sr isotope profile, upper oceanic crust 84, 85 Sr isotope studies 81 analytical methods 132-133 Troodos ophiolite 128 Yaman Kasy 256 Sr partition coefficient 131 Sr ratios dykes 145 pillow lavas 131 secondary minerals 86, 87, 93 transport model 131 Troodos ophiolite 130, 134-135, 136, 137-138, 139, 143, 144, 146 variation with depth 91 Sr/Ca ratios 93 stockworks 82 MORB extrusives 233 sulphide deposits 84 Troodos ophiolite 154, 155, 156, 171 Urals 242 stress modelling 31 strontium, see Sr structural evolution, Kizildag Ophiolite 56 sub-basement, Sr ratios 86 subducting slab, metasomatism 61, 63 subduction, Neo-Tethys 234 subduction zone, oblique 62 submersibles, Nautile 1, 5 sulphide deposits Central Hill 191 clastic 180 Devonian 262 disseminated 191, 212 fault-associated 54 formation depth 255 fossiliferous 260 Kizildag 52 layered 245 lenses 241,242, 256 Lucky Strike segment 2 massive 177, 181,235, 242 metal distribution in 186 mineralization and geochemistry 184 ODP Hole 504B 91 Silurian 263 stacked 181, 188 stockworks 84 Troodos 155 sulphide feeder zones 181, 188 sulphide grains, graded 242 sulphide tolerance, vent communities 283 sulphide toxicity 283 sulphophile faunas 274, 283 sulphur igneous source 255, 257 native 191 sulphur isotopes 93, 250, 255, 256 suprasubduction 63, 65, 128, 155, 212

symbiosis, vent communities 283 symmetry, extrusives 10, 12 TAG 55, 76 major and minor elements 205 ochres 218 rare-earth elements 207, 208 sulphide mound 156, 202, 243 Tahtak6prfi fault 46, 52, 54, 56 talc 110, 116 talus ramps 8, 40 Tauride platform 63 taxonomic characters, vent fossils 266 tear faults, Kizildag 53 tectonic features MAR 5 segment OH3 I0 tectonic stretching 44, 53 magnitude 56 tectonites 33, 36, 46 tectosilicates 110, 116 tellurides 247, 249, 253, 253, 254, 257 Temnocinclus 279 temperatures, black smokers 90 terrigenous processes 229 Tethyan area, ophiolites 219 Tevnia 260, 263 thermal burst, Troodos 63 thermal evolution, crust 149 thermal modelling, fluid fluxes 127 tholeiites 222 thrust faulting 53 thrust sheets, Greece 220, 221,233 thrust stack 222 thrusting, in-sequence 221 Ti depletion, Kizildag 61 Ti/Zr ratios 58 TOBI 1, 2 TOBI images FAMOUS segment 12 Lucky Strike segment 8, 10 Tonga Ridge 219 trace element analysis, Kizildag 57, 60 Trans-Atlantic geotraverse, see TAG transfer faults 56 transform faults 52, 54 transition zone, Sr ratios 84, 85 troctolite 33 Troodos Massif 128 Troodos ophiolite 46, 62, 63, 127 age of 133 alteration pattern 148 Basal Group 136, 143 cumulates 154 dyke-pluton transition 136 fluid fluxes 131 fossils 260 geological map 133, 155 rare-earth element mobility 153 sampling maps 140, 141 serpentinite 154 Sr isotope studies 128 Sr ratios 130 stockworks 154, 155

INDEX sulphide deposits 155 Upper Pillow Lavas 63, 143 tubes, characteristics 261 turbidites, Juan de Fuca Ridge 103, 178, 179 Twin Hills segment 6 U enrichment 203 U/Fe ratios 203, 206, 209 ultramafic/mafic contact 33 ultramafics, fabric 35 ultramylonites 36 umbers 52, 20l, 217 distal deposition 233 plume deposits 218 up-flow zones 119, 157 Upper Pillow Lavas, Troodos 63 upwelling, hydrothermal fluids 129, 138 Uralian Ocean 242 Urals geological map 242 hydrothermal chimneys 241 Uzelga 262 V/Fe ratios 203, 208, 209, 210, 213 valleys, fracture zone 76 Valu Fa Ridge 219 vein minerals, Sr ratios 94, 138 veinlets 86 veins cross-cutting 84 Kizildag plutonics 46 upper crust 82 Urals 242 velocity models, AMC 22, 23 velocity/temperature curves 24 vent communities 259 colonisation 283 endemism 274 faunal origins 282 fossil and modern 267 mosaic origins 285 physiological barriers 282-283 sulphide tolerance 283 symbiosis 283 taxonomic comparisons 267 see also faunas vent sites, fossilization at 264-266 vesicle filling 115, 116 Vesicomya 281 vesicomyids 263, 281 vestimentiferans 260, 263, 264, 266, 267, 274 see also Pogonophora

303 Vetigastropoda 279 volcanic cycles 19 volcanic evolution, three-stage 17 volcanic ridges FAMOUS segment 14 Lucky Strike segment 2, 10 OH3 segment 8 volcanism explosive 6 off-axis 236 volcano-tectonic variations, sketch sections 11 volcanoes conical 9 fissure 31 Lucky Strike segment 2 vugs 86, 87, 185, 191 W-Seamount 31 Waikalasma 278 wall rock alteration 157 water in lower crust and mantle 39 see also fluid, hydrous, seawater etc water-rock ratios 120, 121, 130, 172 waveform inversion procedure 22 waveforms, AMC 20, 21 weathering, TAG mound 203, 209 wehrlite 46 West Antilles arc 229 West Valley 178 white smokers EPR 20 TAG site 205, 210 Woodlark basin 63 worm tubes 243, 255, 260 wurzite 185, 247 xenoliths, ultramafic 35 Yaman Kasy 241,242, 251 footwall 244 fossils 263, 264, 265 geological map and section 243 Yubileinoe 262 Zambales ophiolite 260 zeolites ll0, 116, 136, 142, 147, 165 Zn, in massive sulphides 186, 188 Zuha gossan 205, 212, 213 geochemistry 208