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doi:10.2204/iodp.proc.309312.101.2006

Background and objectives

More than 60% of the Earth is covered by oceanic crust formed at the mid-ocean ridges, all of which formed within the last 200 m.y. The formation of new oceanic crust at ridge axes and reincorporation of aged crust into the mantle or its accretion onto continental margins at subduction zones are perhaps the most fundamental components of the plate tectonic cycle. These processes control the physiography of the Earth and the chemical and thermal evolution of the crust and mantle.

Accretion of oceanic crust at mid-ocean ridges from magmas passively upwelled during partial melting of decompressed mantle peridotite is the dominant process of thermal and chemical transfer from the Earth’s interior to the crust, overlying oceans, and atmosphere. Conduction and hydrothermal advection of heat at constructive plate boundaries through the oceanic lithosphere and as it matures on the ridge flanks are the major mechanisms for heat loss from the interior of the planet. Such hydrothermal interactions influence the chemistry of many elements (e.g., Mg, Sr, U, and K) in the oceans at a similar magnitude to the inputs from rivers draining the continents. Chemical exchange of seawater with the ocean crust leads to major changes in the chemical composition and physical properties of oceanic basement, which, through subduction, influence the composition and heterogeneity of the mantle and the melting and constructional processes occurring in arcs.

Recently, the upper oceanic crust has been shown to contain habitats for microorganisms. Microbes colonize fractures in glassy basaltic rocks, extracting energy and/or nutrients from the glass by dissolving it, leaving behind biomarkers that reveal their former presence. The temperature and depth limits of oceanic basement microbiological activity have yet to be explored, but microbial processes occurring in the submarine deep biosphere may hold the key to the development and survival of life on the Earth and other planets.

Despite the central role that the ocean crust plays in the evolution of our planet, our sampling of in situ oceanic basement is poor, and consequently, our understanding of the fundamental processes involved in the formation and evolution of the oceanic crust remains rudimentary. Samples of basalts, dikes, gabbros, and peridotites have been retrieved by dredging and shallow drill holes from most of the ocean basins, but the geological context of these samples is rarely well established. As such, the nature and variability of the composition and structure of the ocean crust away from transform faults and other tectonic windows remain poorly known. Drilling a complete crustal section has always been a major goal of scientific ocean drilling (Bascom, 1961; Shor, 1985), but achievement of this goal has been impeded by technical difficulties and the time investments required. The distribution of drill holes in intact oceanic crust of different ages and formed at different spreading rates is extremely sparse (Fig. F1) (see Wilson, Teagle, Acton, et al., 2003, for full documentation of basement drilling). Before Integrated Ocean Drilling Program (IODP) Expedition 309, Hole 504B, drilled during the Deep Sea Drilling Project (DSDP) and Ocean Drilling Program (ODP) on the southern flank of the Costa Rica Rift, was the only hole to penetrate a complete sequence of extrusive lavas and partially through the underlying sheeted dike complex (Alt, Kinoshita, Stokking, et al., 1993). The dike/​gabbro boundary has never been drilled, and the nature of the plutonic rocks directly subjacent to the sheeted dike complex is not known, despite this zone being perhaps the most influential in determining the mechanisms of crustal accretion and the geometry of magmatic and hydrothermal interactions. Our poor sampling of ocean crust at different spreading rates and crustal ages and the absence of information on crustal variability compromise our ability to extrapolate observations from specific sites to global descriptions of magmatic accretion processes and hydrothermal exchange in the ocean crust.

Oceanic crust formation and evolution are primary themes for investigation in the Initial Science Plan for the Integrated Ocean Drilling Program (International Working Group, 2001) and other major science priority submissions (e.g., Conference on Multiple Platform Exploration of the Ocean [COMPLEX], Pisias and Delaney, 1999; Ocean Drilling Program Geochemistry Futures Workshop, Murray et al., 2002). These documents and others specifically related to the study of the oceanic lithosphere (Second Conference on Scientific Ocean Drilling [COSOD II]; ODP Long Range Plan, Ocean Drilling Program, 1996; ODP–International Cooperation in Ridge-Crest Studies–International Association of Volcanology and Chemistry of the Earth’s Interior workshop; 4D-Architecture of the Ocean Crust Program Planning Group) reemphasize deep drilling to obtain complete sections of the ocean crust as a priority and note that the deep drilling capabilities of riserless technology have yet to be fully utilized.

Offset drilling strategies, where deeper portions of the ocean crust are sampled by drilling in tectonic windows, have recently been high priorities for ocean drilling (COSOD II, 1987; Ocean Drilling Program, 1996). Drilling at several sites has provided a wealth of new data and understanding of gabbros and peridotites from the lower crust and serpentinized upper mantle (e.g., Hess Deep: Gillis, Mével, Allan, et al., 1993; Kane Fracture Zone area: Cannat, Karson, Miller, et al., 1995; Southwest Indian Ridge: Dick, Natland, Miller, et al., 1999; 14°–16°N Mid-Atlantic Ridge: Kelemen, Kikawa, Miller, et al., 2004; Atlantis Massif: Expedition 304 Scientists, 2005, Expedition 305 Scientists, 2005). However, serious problems still exist with drilling tectonized rocks with little sediment blanket or without erosional removal of fractured material, and it is also commonly difficult to relate drilled sections to regional geology (e.g., Gillis, Mével, Allan, et al., 1993). Perhaps the most successful deep drilling of the oceanic basement has occurred in crust formed at slow to ultra-slow spreading rates on plutonic massifs exposed near ridge-transform fault intersections (Holes 735B and U1309D). These holes boast deep penetration at high rates of recovery (Dick, Natland, Miller, et al., 1999; Expedition 304 Scientists, 2005; Expedition 305 Scientists, 2005) but are probably not representative of the majority of ocean crust that formed at faster spreading rates and in the middle of ridge segments.

Composite sections are not substitutes for deep in situ penetrations, and drilling deep holes to obtain complete upper crustal sections continues to be a primary challenge for scientific ocean drilling (Dick and Mével, 1996; Murray et al., 2002). Unfortunately, there are no on-land alternatives to drilling in the oceans. Although ophiolites, ancient slices of ocean crust now preserved on land, provided much of the early inspiration for ocean crust studies, the classic outcrops of Semail ophiolite (Oman) and Troodos massif (Cyprus) formed in suprasubduction zone settings reference (e.g., Pearce and Cann, 1971; Pearce et al., 1981; Rautenschlein et al., 1985) and their different magma and volatile chemistries compromise their applicability for understanding processes in the major ocean basins. Macquarie Island (Varne et al., 2000), uplifted along the Australian/​Pacific plate boundary ~1000 km south of New Zealand, may be the only outcrop of subaerially exposed ocean crust formed at a mid-ocean ridge, but the island is complexly faulted and is an environmentally sensitive United Nations Educational, Scientific, and Cultural Organization World Heritage site from which drilling, even for scientific purposes, is prohibited.

IODP Expedition 309/312 builds on the success of ODP Leg 206, during which Hole 1256D was established, to drill a complete section of the upper oceanic crust. Hole 1256D in the Guatemala Basin is hosted by 15 m.y. old crust that formed at the equatorial East Pacific Rise (EPR) during a sustained period of superfast seafloor spreading (Fig. F2) (Wilson, Teagle, Acton, et al., 2003). It is the first basement borehole prepared with the infrastructure desirable for drilling a moderately deep hole into the oceanic crust (~1.5–2 km). Following preliminary coring to document the sedimentary overburden at Site 1256, a reentry cone supported by 20 inch casing and large-diameter (16 inch) casing all the way through the sediment cover and cemented 19 m into basement was installed in Hole 1256D during Leg 206. The cone and casing facilitates multiple reentries and helps to maintain hole stability essential for deepening Hole 1256D down through the dikes and into the gabbros. The large-diameter casing leaves open the possibility that one or two more casing strings could be installed in Hole 1256D should future expeditions need to isolate unstable portions of the hole. During Leg 206, the upper 502 m of the igneous crust was cored with moderate to high recovery (average = ~48%), and the uppermost crust at this site comprises a sequence of massive flows and thinner sheet flows with subordinate pillow basalt and breccias. The sequence has normal mid-ocean-ridge basalt (N-MORB) composition similar to modern EPR basalts and is slightly to moderately altered. It was extruded over sufficient time to record stable geomagnetic field directions and capture transitional directions in the upper units as the geomagnetic field reversed. Importantly, operations during Leg 206 in Hole 1256D concluded with the hole clean of debris and in excellent condition.

The goal of the IODP Superfast Spreading Rate Crust mission (Expedition 309/312) was to core through the remaining extrusive rocks and the underlying sheeted dike complex and into the upper gabbros. This continuous section of in situ oceanic crust generated at a superfast spreading rate in the eastern Pacific will

• Provide the first sampling of a complete section of ocean crust from the extrusive rocks and dikes and into the gabbros. This will confirm whether ocean crust formed at a superfast spreading rate conforms to a “Penrose” ophiolite stratigraphy.
• Confirm the nature of high-level axial magma chambers.
• Define the relationship between magma chambers and their overlying lavas and the interactions between magmatic, hydrothermal, and tectonic processes.
• Provide in situ calibration of seismic velocity and magnetic measurements made from surface ships so that these regional-scale geophysical measurements can be related to geology.

Rationale for site selection and location criteria for deep drilling

The key to proposing the Superfast Spreading Rate Crust campaign (Leg 206 and Expedition 309/312) was to identify a style of crustal accretion where the extrusive lavas and dikes overlying the gabbros were predicted to be relatively thin, thus increasing the likelihood of penetrating through the complete upper crustal section in the fewest drilling days. The recognition that crust formed at a superfast spreading rate is a compelling target for deep drilling follows the observation that there is an inverse relationship between the depth of axial low-velocity zones imaged by seismic experiments, interpreted to be melt lenses, and spreading rate (Purdy et al., 1992) (Fig. F3). Even allowing for an additional thickness of lavas that flowed from the ridge axis to cover the immediate flanks, the uppermost gabbros should be at relatively shallow depths in superfast-spreading crust. The predicted depth to gabbros at Site 504 on the south flank of the intermediate-spreading Costa Rica Rift is >2.5 km, whereas the depth to the axial low-velocity zones at typical fast spreading rate (~80–150 mm/y) crust on the EPR is ~1–2 km. The estimated depth to an axial melt lens for ocean crust formed at a superfast spreading rate is ~700–1000 m, and the anticipated depth to the gabbros for Site 1256 is ~1000–1300 m, allowing for a reasonable thickness (~300 m) of near-axial lava flows.

Although perhaps only 20% of the global ridge axis is separating at fast spreading rates (>80 mm/y full rate), ~50% of the present-day ocean crust and ~30% of the Earth’s total surface was produced by this pace of ocean spreading. At least in terms of seismic structure (Raitt, 1963; Menard, 1964), crust formed at fast spreading rates is relatively simple and uniform. Deep drilling at Site 1256 will characterize one end-member style of mid-ocean-ridge accretion, and the successful deep sampling of such crust in a single location can reasonably be extrapolated to describe a significant portion of the Earth’s surface.

Site selection

A recent reconsideration of magnetic anomalies at the southern end of the Pacific/​Cocos plate boundary has identified crust formed at a full spreading rate of ~220 mm/y from 20 to 11 Ma (Wilson, 1996) (Fig. F2). This is significantly faster than the present fastest spreading rate (~145 mm/y) for crust forming at ~30°S on the EPR. From this region created by superfast spreading, a single drill site in the Guatemala Basin, initially designated GUATB-03C and now known as Site 1256, was selected on young, 15 m.y. old ocean crust. The details of site survey operations and the reasons for the selection of this particular site are outlined in detail in Wilson, Teagle, Acton, et al. (2003).

In addition to the shallow depth to gabbros predicted from formation at a superfast spreading rate, Site 1256 has a number of specific attributes that indicate that this site provides an excellent opportunity to sample a complete section of upper oceanic crust. Site 1256 formed at an equatorial latitude (Fig. F4), and high equatorial productivity resulted in robust sedimentation rates (>30 m/m.y.) and the rapid burial of the young basement. The thick sediment blanket enabled installation of a reentry cone with 20 inch casing that forms the foundation for deployment of the second 16 inch diameter casing string that was cemented 19 m into the uppermost basement. At 15 Ma, Site 1256 is significantly older than the crust at Hole 504B (6.9 Ma), and lower temperatures are anticipated at midlevels of the crust. As such, high basement temperatures, which can preclude drilling operations, should not be reached until gabbroic rocks have been penetrated. Logistically, Site 1256 has a number of advantages. It is ~3.5 days steaming from the Panama Canal, and the short transit time allows for maximum time on site during drilling expeditions. As transfer between the Pacific and Atlantic Oceans is common because of the scheduling demands of scientific drilling, close proximity to the Panama Canal has allowed the timely rescheduling of return visits to the site.

Geological setting of Hole 1256D

Site 1256 (6°44.1′N, 91°56.1′W) lies in 3635 m of water in the Guatemala Basin on Cocos plate crust formed ~15 m.y. ago on the eastern flank of the EPR (Figs. F2, F4). The depth of the site is close to that predicted from bathymetry models of plate cooling (e.g., Parsons and Sclater, 1977). The site sits astride the magnetic Anomaly 5Bn–5Br magnetic polarity transition (Fig. F5A). This crust accreted at a superfast spreading rate (~220 mm/y full rate; Wilson, 1996) and lies ~1150 km east of the present crest of the EPR and ~530 km north of Cocos Ridge. The site formed on a ridge segment at least 400 km in length, ~100 km north of the ridge-ridge-ridge triple junction between the Cocos, Pacific, and Nazca plates (Fig. F4). This location was initially at an equatorial latitude within the equatorial high-productivity zone and endured high sedimentation rates (>30 m/m.y.; e.g., Farrell et al., 1995; Wilson, Teagle, Acton, et al., 2003). Sediment thickness in the region is between 200 and 300 m and is 250 m at Site 1256 (Wilson, Teagle, Acton, et al., 2003).

Site 1256 has a seismic structure reminiscent of typical Pacific off-axis seafloor (Fig. F6). Upper Layer 2 velocities are 4.5–5 km/s and the Layer 2–3 transition is between ~1200 and 1500 meters subbasement (msb) (Fig. F7). Total crustal thickness at Site 1256 is estimated at ~5–5.5 km. Further to the northeast of Site 1256 (15–20 km), a trail of ~500 m high circular seamounts rise a few hundred meters above the sediment blanket (Fig. F5B).

Using the site survey multichannel seismic (MCS) data (Wilson et al., 2003), we have constructed a geological sketch map of the uppermost basement in the GUATB-03 survey region (Fig. F8). The bathymetry in the GUATB-03 survey area is generally subdued, and Site 1256 sits atop a region of smooth basement topography (<10 m relief). However, elsewhere in the region the top of basement shows a number of offsets along northwest-striking normal faults, and an abyssal hill relief of up to 100 m is apparent in the southwest part of the area. Relief to the northeast is lower and less organized. In the northeast sector of the GUAT-3B region, there is evidence for a basement thrust fault with a strike approximately orthogonal to the regional fabric (Wilson et al., 2003; Hallenborg et al., 2003). This feature dips gently to the northwest (~15°) and is clearly discernible to a depth of ~1.3 km on seismic Line EW9903-28 (Wilson et al., 2003), but the feature is less pronounced on seismic Line EW9903-27, indicating that the offset on the thrust decreases to the southwest.

Additional processing (A. Harding, unpubl. data) of ocean bottom seismometer recordings indicates that there is discernible variation in the average seismic velocity (~4.54–4.88 km/s) of the uppermost (~100 m) basement and that there is regional coherence in the velocity variations (Fig. F8A). Two principal features are apparent: a 5–10 km wide zone of relatively high upper basement velocities (>4.82 km/s) that can be traced ~20 km to the edge of data coverage southeast of Site 1256 and a relatively low velocity (4.66–4.54 km/s) bull’s-eye centered around the crossing point of seismic Lines EW9903-21 and 25.

The uppermost basement at Site 1256 is capped by a massive lava flow >74 m thick. This flow is relatively unfractured, with shipboard physical properties measurements on discrete samples indicating V_P > 5.5 km/s (Wilson, Teagle, Acton, et al., 2003). As such, it is likely that the area of relatively high uppermost basement seismic velocities delineates the extent of the massive flow penetrated in Holes 1256C and 1256D. Assuming an average thickness of 40 m, this would conservatively suggest an eruption volume >3 km³, plausibly >10 km³. This is extremely large when compared to the size of mid-ocean-ridge axial low-velocity zones that are thought to be high-level melt lenses and which typically have volumes ~0.05–0.15 km³/km of ridge axis and generally appear to be only partially molten (Singh et al., 1998).

Sheet flows (<3 m thick) and massive flows (>3 m) make up most of the lava stratigraphy cored at Site 1256, and such lava morphologies dominate crust formed at fast spreading rates, away from segment tips (e.g., White et al., 2000, 2002). Subordinate pillow lavas are present in Hole 1256D, and because of the large number of fractures and pillow interstices, seismic velocities are generally lower than more massive lava flows. We speculate that the bull’s eye of relatively low seismic velocities is a thick pile of dominantly pillowed lava flows.

Scientific objectives of Expedition 309/312

During Expedition 309/312, drilling of a continuous section through volcanic basement and the underlying sheeted dike complex down into the uppermost plutonic rocks in Hole 1256D continued. The cores recovered from and the wireline measurements made in Hole 1256D will provide unique information to address the following specific scientific objectives (Table T1):

1. Test the prediction, from the correlation of spreading rate with decreasing depth to the axial low-velocity zones (e.g., Purdy et al., 1992), that gabbros representing the crystallized melt lens will be encountered at 1000–1300 msb at Site 1256.

The transition from sheeted dikes to gabbros has never been drilled, and this remains an important objective in achieving a complete or even composite oceanic crustal section. The dike–gabbro transition and the uppermost plutonic rocks are assumed to be the frozen axial melt lens and the fossil thermal boundary layer between magma chambers and vigorous hydrothermal circulation. Detailed knowledge of the dike–gabbro transition zone is critical to discerning the mechanisms of crustal accretion. The textures and chemistries of the uppermost gabbros are presently unknown but are central to understanding crustal construction. At present, we lack samples that link gabbroic rocks to the overlying lavas, leading to the following questions:

• What is the geological nature of the low-velocity zones imaged by MCS reflection studies at the axes of mid-ocean ridges?
• Is the melt lens imaged at mid-ocean ridges made of cumulate rocks from which magmas are expelled to form the dikes and lavas, which then subside to form the lower crust? Or are the uppermost gabbros coarse-grained chemical equivalents of the dikes and extrusive rocks frozen at the base of the sheeted dikes?
• Does most of the crustal accretion occur at deeper levels through the intrusion of multiple thin sills?
• What are the cooling rates of magma chambers?

These questions can be answered through petrological and geochemical studies of gabbros (e.g., Natland and Dick, 1996; Kelemen et al., 1997; Manning et al., 2000; MacLeod and Yaoancq, 2000; Coogan et al., 2002a, 2002b) and the overlying lavas and their mineral constituents.

2. Determine the lithology and structure of the upper oceanic crust from a superfast spreading rate end-member.

Some basic observations regarding the architecture of ocean crust, including the lithology, geochemistry, and thicknesses of the volcanic and sheeted dike sections and how these vary with spreading rate or tectonic setting, are not well known. Karson (2002) provides estimates of the thicknesses of lavas and sheeted dikes from crust generated at fast and intermediate spreading rates (600–900 m lavas and 300–1000 m dikes at Hess Deep; 500–1300 m lavas and 500 to >1000 m dikes in Hole 504B and at the Blanco Fracture Zone). With the exception of the incomplete section in Hole 504B, these estimates are based on observations of highly disrupted exposures, where structural complexities and the uniqueness of the geological environments indicate that such estimates should be treated with caution. The results of Expedition 309/312 will determine the thicknesses of these upper crustal units at Site 1256 and document the styles of deformation and magmatic accretion. Studies of tectonic exposures of oceanic crust suggest that intense brittle deformation, faulting and distributed zones of fracturing, and large amounts of dike rotation are common within sheeted dike complexes in crust formed at fast and intermediate spreading rates (Karson, 2002; Karson et al., 2002; Stewart et al., 2005). It is difficult, however, to separate primary mid-ocean-ridge geometries from deformation related to the unroofing of these tectonic windows. In contrast, large blocks of the sheeted dike complex in the Semail ophiolite in Oman exhibit little of such faulting or distributed fracturing (Umino et al., 2003). Seismic profiles of the Site 1256 region show well-developed subhorizontal reflectors to ~800–900 msb (Fig. F6), providing little evidence for rotation of the upper crust in this region.

Drilling the sheeted dike complex at Site 1256 will enable evaluation of whether such faulting and fracturing observed in tectonic exposures are representative of oceanic crust or merely related to their complex tectonic settings. Most dikes in sheeted dike complexes in tectonic exposures of crust generated at intermediate and fast spreading rates and in Hole 504B in intermediate rate crust generally dip away from the spreading axis, suggesting tectonic rotation of crustal blocks (Karson, 2002). Do such rotations also occur in crust generated at superfast spreading rates, and are they similar, or is the crust less tectonically disrupted? A single drill hole may not conclusively answer this question but should provide important constraints.

3. Correlate and calibrate seismic and magnetic imaging of the crustal structure with basic geological observations.

Ground-truthing regional geophysical techniques such as seismic and magnetic imaging is a key goal of the IODP Initial Science Plan and related documents (e.g., COMPLEX). A fundamental question we will address in this experiment is how velocity changes within seismic Layer 2 and whether the Layer 2–Layer 3 transition relates to physical, lithological, structural, and/or alteration variations in the volcanic rocks, dikes, and gabbros. At Site 504, in crust generated at an intermediate spreading rate ridge, the Layer 2–Layer 3 transition lies within the 1 km thick sheeted dike complex and coincides with a metamorphic change (Detrick et al., 1994; Alt et al., 1996), but it is unknown whether the results from Hole 504B are representative of ocean crust in general or of crust generated at different spreading rates. Is the depth to gabbros shallower in crust generated at a superfast spreading rate, as predicted, and what are the relative thicknesses of volcanic and dike sections compared with crust constructed at slow or intermediate spreading rates?

Marine magnetic anomalies are one of the key observations that led to the development of plate tectonic theory, through recognition that the ocean crust records the changing polarity of the Earth’s magnetic field through time (Vine and Matthews, 1963). It is generally assumed that micrometer-sized grains of titanomagnetite within the erupted basalts are the principal recorders of marine magnetic anomalies. However, recent studies of tectonically exhumed lower crustal rocks and serpentinized upper mantle indicate that these deeper rocks may also be a significant source of the magnetic stripes (Hosford et al., 2003). Coring a complete section through the sheeted dike complex will allow evaluation of the contribution of these rocks to marine magnetic anomalies. Whether these deeper rocks have a significant influence on the magnetic field in undisrupted crust is unknown, as is the extent of secondary magnetite growth in gabbros and mantle assemblages away from transform faults. Sampling the plutonic layers of the crust will test the Vine-Matthews hypothesis by characterizing the magnetic properties of gabbros through drilling normal ocean crust on a well-defined magnetic stripe, away from transform faults.

4. Investigate the interactions between magmatic and alteration processes, including the relationships between extrusive volcanic rocks, sheeted dikes, and underlying gabbroic rocks.

Little information presently exists on the heterogeneity of hydrothermal alteration in the upper crust or the variability of associated thermal, fluid, and chemical fluxes. How these phenomena vary at similar and different spreading rates is unknown. Metamorphic assemblages and analyses of secondary minerals in material recovered by deep drilling can provide limits on the amount of heat removed by hydrothermal systems and place important constraints on the geometry of magmatic accretion and the thermal history of both the upper and lower crust (e.g., Manning et al., 2000; MacLeod and Yaoancq, 2000; Coogan et al., 2002a, 2002b). Fluid flow paths, the extent of alteration, and the nature of deep subsurface reaction and shallower mixing zones are all critical components of our understanding of hydrothermal processes that can only be tackled by drilling. These problems can be addressed by examining the “stratigraphy” and relative chronology of alteration within the extrusive lavas and dikes, by determining whether disseminated sulfide mineralization resulting from fluid mixing and a large step in thermal conditions is present at the volcanic–dike transition (as in Hole 504B and many ophiolites), and by evaluating the grade and intensity of alteration in the lower dikes and upper gabbros. The lowermost dikes and upper gabbros have been identified as the conductive boundary layer between the magma chambers and the axial high-temperature hydrothermal systems, as well as the subsurface reaction zone where downwelling fluids acquire black-smoker chemistry (Alt, 1995; Alt et al., 1996; Vanko and Laverne, 1998; Gillis et al., 2001). However, extensive regions of this style of alteration or zones of focused discharge are poorly known, and information from ophiolites may not be applicable to in situ ocean crust (Richardson et al., 1987; Schiffman and Smith, 1988; Bickle and Teagle, 1992; Gillis and Roberts, 1999). Drilling beyond the boundary between the lower dikes and upper gabbros will help trace recharge fluid compositions, estimate hydrothermal fluid fluxes (e.g., Teagle et al., 1998, 2003; Laverne et al., 2001; Gillis et al., 2005), and integrate the thermal requirements of hydrothermal alteration in sheeted dikes and underlying gabbros with the magmatic processes in the melt lens. Detailed logging of cores combined with geochemical analyses will enable determination of geochemical budgets for hydrothermal alteration (e.g., Alt et al., 1996; Alt and Teagle, 1999, 2000; Bach et al., 2003). Is there a balance between the effects of low-temperature alteration of lavas versus high-temperature hydrothermal alteration of dikes and gabbros? This is a critical check on global budgets for many elements (Mg, K, ⁸⁷Sr, U, and ¹⁸O) presently estimated from vent fluid chemistries, riverine inputs, and thermal models (e.g., review of Elderfield and Schultz, 1996).

The discovery of microorganisms that colonize and extract energy from volcanic glass in the upper oceanic crust has added new dimensions to seafloor alteration studies. Microbial alteration of volcanic glass has been shown to decrease with basement depth at other sites (Furnes and Staudigel, 1999). The temperature and depth limits to subbasement microbiological activity remain unknown but can be investigated by deep sampling and study of microbial alteration textures, chemical and isotopic indicators, and molecular microbiology (e.g., Blake et al., 2001; Alt et al., 2003; Banerjee and Muehlenbachs, 2003).

Principal results of ODP Leg 206

The major objectives of Leg 206 were to establish a cased reentry hole that is open for future drilling and to achieve a target penetration in excess of 500 msb. Before basement drilling was initiated, a series of holes was drilled to thoroughly characterize the sediments and magneto- and biostratigraphy of the sedimentary overburden and to determine the casing depth into basement for the main hole. Four holes were drilled during Leg 206, with Holes 1256A, 1256B, and 1256C recovering a near-complete record of the 250 m thick sedimentary overburden. Pilot Hole 1256C penetrated 88.5 m into basement, with Hole 1256D being the cased reentry hole with a large reentry cone supported by 95 m of 20 inch casing and 269.5 m of 16 inch casing cemented into the uppermost basement. The total depth of penetration of Hole 1256D during Leg 206 was 752 mbsf, including 502 m drilled into basement. Recovery of igneous rocks was good, and excellent in places, with average recovery rates of 61.3% and 47.8% in Holes 1256C and 1256D, respectively (Wilson, Teagle, Acton, et al., 2003).

The sedimentary overburden is divided into two units: Unit I (0–40.6 mbsf) is clay rich, with a few carbonate-rich layers; Unit II (40.6–250.7 mbsf) is predominantly biogenic carbonate. The interval from 111 to 115 mbsf is rich in biogenic silica, which forms a distinct diatom mat, deposited at ~10.8 Ma. Chert nodules are a common feature below 111 mbsf, and red-brown iron oxide–rich silicified sediments that may be recrystallized metalliferous sediments are present directly over the basement (within ~1 m). The primary control on the interstitial water chemistry at Site 1256 is diffusion between seawater and basement fluids, with a continuous chert bed at 158 mbsf providing a low-diffusivity barrier. Calcareous microfossil biostratigraphy is in good agreement with magnetostratigraphy. Calculated sedimentation rates vary from ~6 to 36 m/m.y. and decrease with time as the site moved away from the high-productivity zone near the paleoequator and the cooling lithospheric plate subsided (Wilson, Teagle, Acton, et al., 2003).

Approximately 60% of the igneous basement in Holes 1256C and 1256D consists of thin (tens of centimeters to <3 m) basaltic sheet flows separated by chilled margins (Fig. F9). Massive flows (>3 m thick) are the second most common rock type, including the thick ponded flow near the top of the holes. Minor intervals of pillow lavas (~20 m) and hyaloclastite (a few meters) and a single dike were recovered in Hole 1256D. The low proportion of pillow lavas (<10%) indicates rapid lava emplacement on low topographic relief, consistent with thermal model predictions of <1 km vertical thickness of the dike zone.

The uppermost lavas, sampled only in Hole 1256C because of setting casing in Hole 1256D, are composed of thin basaltic sheet flows a few tens of centimeters to ~3 m thick, separated by chilled margins and containing rare intervals of recrystallized sediment. Basement Units 1256C-18 and 1256D-1 each consist of a single cooling unit of cryptocrystalline to fine-grained basalt, interpreted to be a ponded lava flow and serving as a clear marker unit for correlation of the igneous stratigraphy between holes. A total of 32 m of this unit was cored in Hole 1256C, of which 29 m was recovered. This ponded flow is much thicker 30 m away in Hole 1256D, where it has a minimum thickness of 74.2 m, indicating steep paleotopography. The groundmass of the interior of the flow is fine grained but is deformed and thermally metamorphosed ~1.5 m from its base.

The transition from axial eruptions to lavas that flowed out onto the ridge flanks was not determined during Leg 206. However, it was recognized that the thickness of the massive ponded flow requires significant basement relief in order to pool the lava, and this would only be developed significantly off axis (5–10 km).

The remainder of the section in Hole 1256D (with the exception of Units 1256D-3, 4a, 4c, 8c, 16d, and 21) consists of sheet flows tens of centimeters to ~3 m thick with uncommon massive flows 3.5–16 m thick. These sheet and massive flows are aphyric to sparsely phyric, cryptocrystalline to microcrystalline basalt and are distinguished by common chilled margins with fresh or altered glass.

One ~20 m thick interval of aphyric to sparsely phyric cryptocrystalline pillow basalt with glassy chilled margins was recovered from near the top of the section (Unit 1256D-3), as well as two 1.0–1.7 m thick hyaloclastite intervals (Units 1256D-4c and 21). Also recovered was a 0.3 m thick interval of volcanic breccia composed of angular fragments of cryptocrystalline basalt embedded in a matrix of altered glass (Unit 1256D-4a).

The basalts show large variations in grain size and textures from holohyaline in the outermost chilled margins of lava flows to the coarser intergranular textures in the lava pond. The basaltic lavas are dominantly aphyric to sparsely phyric, but where phenocrysts are present, olivine is the dominant phase with subordinate plagioclase, minor clinopyroxene, and rare spinel. Measurements of petrographically fresh samples revealed general downhole variations with Mg# (= Mg/[Mg + Fe]), Cr, Ni, and Ca/Al ratios broadly increasing with depth, whereas TiO₂, Fe₂O₃, Zr, Y, Nb, V, and Sr broadly decrease with depth, although smaller-scale variations are superimposed on these trends (Fig. F9). All Leg 206 lavas from Site 1256 plot in the N-MORB field on a Zr-Y-Nb ternary diagram.

In the lavas directly below the large massive flow, there is a sharp increase in Mg# accompanied by an increase in incompatible element concentrations (Fig. F9). The combination of high Mg# and high incompatible element concentrations argues against differentiation as the cause of the enrichments and suggests that there is variation in the primitive magma composition.

Rocks throughout Holes 1256C and 1256D exhibit a dark gray background alteration, where the rocks are slightly to moderately altered and olivine is replaced and pore spaces are filled by saponite and minor pyrite as the result of low-temperature seawater interaction at low cumulative seawater/rock ratios. Vein-related alteration manifests as different-colored alteration halos along veins. Black halos contain celadonite and result from the reaction of young ocean crust (<1–2 Ma) with distal upwelling low-temperature hydrothermal fluids enriched in iron, silica, and alkalis (Edmond et al., 1979; see summary in Alt, 1999). Iron oxyhydroxide–rich brown mixed halos are later features, which formed by circulation of oxidizing seawater. Brown halos have a similar origin and formed along fractures that were not bordered by previously formed black halos. This vein-related alteration occurs irregularly throughout Hole 1256D below the massive Unit 1256D-1 but is concentrated in two zones of greater permeability and, consequently, increased fluid flow, at 350–450 and 635–750 mbsf. The appearance of albite and saponite partially replacing plagioclase below 625 mbsf indicates a change in alteration conditions. This change may result in part because of slightly higher temperatures at depth as the lava/dike boundary is approached or from interaction with more evolved fluid compositions (e.g., decreased K/Na and elevated silica).

When compared with other basement sites, Hole 1256D (Fig. F10) contains far fewer brown, mixed, and black alteration halos. The abundance of carbonate veins in Hole 1256D is also lower than at many other sites. Site 1256 is, however, quite similar to Site 801, also in crust generated at a fast-spreading ridge, albeit 170 m.y. ago. One important feature is the lack of any oxidation gradient with depth in Hole 1256D, in contrast to the stepwise disappearance of iron oxyhydroxide and celadonite in Hole 504B and the general downward decrease in seawater effects at Sites 417 and 418. In contrast, alteration appears to have been concentrated into different zones that may be related to the architecture of the basement, such as lava morphology, distribution of breccia and fracturing, and the influence of these on porosity and permeability. Clearly, there is greater variation in the processes of alteration occurring in the oceanic crust than is recorded at Sites 504, 417, and 418, which have served as “reference” sections to date. This illustrates the point that models for the formation and alteration of oceanic basement based on crust formed at slow and intermediate spreading rates cannot automatically be applied to crust generated at fast spreading rates.

Structures in the basement at Site 1256 include both igneous features and postmagmatic deformation. Igneous structures such as flow layering, preferred phenocryst orientation, or fine layering delineated by coalesced spherulites or vesicles are observed near the chilled margins of sheet flows but are best developed near the upper and basal margin of the massive ponded flow (Units 1256C-18 and 1256D-1). The massive ponded flow exhibits other features not observed in the rest of the hole, such as the folding of flow layering and shear-related structures, highlighting the complex internal dynamics occurring during the emplacement and cooling of this large igneous body. Folds at the top of the ponded lava have gently dipping axial planes, whereas such features become steeper toward the bottom. Shear indicators, such as pull-aparts and tension gashes, now filled with late-stage magma, are more common toward the base of the ponded lava.

Brittle deformation is common throughout the upper crust sampled by Holes 1256C and 1256D and includes veins with various morphologies, shear veins, joints, and breccias. There is no systematic variation in the structural attitude (true dip) with depth, and this probably reflects the influence of other factors, such as grain size or lava morphology, rather than regional tectonics or the local stress field. Aphyric basalt in sheet flows exhibits a more irregular fracture pattern than coarser grained lavas. Shear veins indicate both normal and reverse senses of shear, suggesting the occurrence of some, although probably local, compressional components. Shear veins are most common in the massive ponded lava (Units 1256C-18 and 1256D-1) and in sheet flows from Cores 206-1256D-27R through 43R, where the geometries of the infilling fibers indicate reverse senses of shear. Brecciated rocks of different styles occur throughout the cores but are most common in the sheet flows. Textural features indicate that most breccias formed either by reworking of lava tops (both chilled and cryptocrystalline basalt) or by the fracturing of rock assisted by relatively high fluid pressures.

Basalt samples from Site 1256 show a strong tendency to have been partially or fully remagnetized during drilling, much more so than for most other DSDP and ODP sites. However, in many cases, a preoverprint component can be discerned, if not always measured accurately, with the shipboard equipment. For Hole 1256D, most samples from igneous Units 1256D-3 through 8a and 14 through 26 demagnetize to a shallow inclination, as expected for the equatorial paleolatitude. For Hole 1256C, all samples have steep inclinations and most are dominated by overprint, but a few samples from Units 1256C-3, 7, 18c, 18h, and 22 show evidence for a stable, steep component distinct from the overprint. The steep inclination may reflect eruption during the magnetic polarity transition between Chrons 5Br and 5Bn, which would imply transport of these uppermost lavas >5 km from the ridge axis.

The basement rocks cored during Leg 206 at Site 1256 show little variation in physical properties with depth. The rocks in and above the massive ponded flow (Units 1256C-1 through 18 and 1256D-1; ~276–350 mbsf) have an average bulk density of 2.89 ± 0.03 g/cm³, which is not significantly different from the basalts below when considered together (2.8 ± 0.1 g/cm³). However, there is a significant decrease in average density (2.7 ± 0.1 g/cm³) for the lava flows immediately below Unit 1256D-1 from 350 to 451 mbsf (Units 1256D-2 through 8a). The average porosity of the basement rocks drilled during Leg 206 is 5% ± 3% within a total range of 2%–19%. The average thermal conductivity of basalts from Hole 1256D is 1.8 ± 0.1 W/(m·K). V_P of discrete basalt samples from Leg 206 varies from 4.2 to 6.2 km/s (average = 5.4 ± 0.1 km/s). A notable exception to the uniform velocity structure of the upper 500 m of basement occurs immediately below the massive ponded lava (Units 1256D-2 through 4c; 350–400 mbsf). Here, average velocity decreases 0.6 km/s to 4.8 km/s and variability increases from ±0.1 to ±0.3 km/s. The ponded lavas have a slightly higher discrete sample V_P (5.5 ± 0.1 km/s) than most of the rocks below, which have an average V_P of 5.4 ± 0.1 km/s. Magnetic susceptibility varies from ~0 to 10,000 × 10^–5 SI in the upper 500 m of Hole 1256D basement. The ponded lavas have an average magnetic susceptibility of 5100 × 10^–5 ± 900 × 10^–5 SI, and below 350.3–752 mbsf, magnetic susceptibility values increase systematically with depth, from ~1000 × 10^–5 to 5000 × 10^–5 SI, with an average magnetic susceptibility of 3000 × 10^–5 ± 1800 × 10^–5 SI. Natural gamma ray measurements were rarely above background in the basement rocks from Leg 206, with the exception of a potassium-rich zone (Unit 1256C-18; 294–308 mbsf) in the massive ponded lava.

A full suite of logging tools was run in Hole 1256D following the completion of coring operations. The tools utilized, in order of deployment, were the triple combination (triple combo) tool string, the Formation MicroScanner (FMS)-sonic tool string, the Ultrasonic Borehole Imager (UBI), and the Well Seismic Tool (WST). Logging showed that Hole 1256D was in excellent condition with no constrictions or ledges. Caliper readings from both the triple combo and the FMS-sonic tool strings show the borehole diameter to be mostly between 11 and 14 inches, with only four short intervals >16 inches. The downhole measurements and images recorded show a large amount of variation, reflecting the massive units, lava flows, pillow lavas, and hyaloclastites recovered in Hole 1256D (Fig. F9).

Predictions of depth to gabbros

The recognition that there is a relationship between spreading rate and the depth to axial low-velocity zones, imaged by MCS experiments across the axes of mid-ocean ridges and thought to be axial melt lenses, is fundamental to the pursuit of the Superfast Spreading Rate Crust mission. Extrapolation of the depth to the axial low-velocity zones versus spreading rate relationship (Fig. F3) (Purdy et al., 1992) to a superfast spreading rate akin to that occurring during the accretion of crust at Site 1256 ~15 m.y. ago on the EPR (200–220 mm/y; Wilson, 1996) indicates that the melt lens would have been located at depths between ~725 and 1000 m beneath the axis. As the new plate cools and moves away from the ridge axis, it will become buried by lavas that flow short distances down the ridge slope (~1–2 km), as well as by larger lava bodies that flow significant distances (~5–10 km) off axis, such as the >74 m thick massive lava pond (Units 1256C-18 and 1256D-1) that formed the upper crust at Site 1256 or similar features recognized on the modern EPR (e.g., Macdonald et al., 1989).

In planning Expedition 309/312, the depth to gabbros was estimated by assuming that the total thickness of near-axis and off-axis lavas was between 100 and 300 m, giving a total estimated depth to gabbros of between 825 and 1300 msb (Fig. F11; Table T2). Our selection of 100–300 m of lava flows comes from a number of lines of evidence. MCS experiments on the EPR estimate Layer 2A thicknesses of ~300 m in the near-axis region (Hooft et al., 1996; Carbotte et al., 1997a), although the geological nature of this seismic layer remains poorly understood. Stronger evidence comes from petrologic descriptions of the rocks from the uppermost basement at Site 1256 drilled during Leg 206. The very uppermost basement, designated during Expedition 309 as the "lava pond" (Units 1256C-1 through 18 and 1256D-1; 250–350.3 mbsf), comprises thin basaltic sheet flows a few tens of centimeters to ~3 m thick (Units 1256C-1 through 17) overlying a massive ponded flow (Units 1256C-18 and 1256D-1) of ~30 to ~74 m of fine-grained basalt in Holes 1256C and 1256D, respectively (Wilson, Teagle, Acton, et al., 2003). Although the massive flow is much thicker in Hole 1256D than in 1256C, it is interpreted as a single lava body whose interior was liquid at the same time in both locations. The dramatic increase in thickness over 30 m of lateral distance and a total thickness in excess of 74 m required at least this much paleotopography to pool the lava. On fast-spreading ridges, such topography does not normally develop until ~5–10 km from the axis (e.g., Macdonald et al., 1989), and we suspect that these magmas flowed a significant distance off axis before ponding in a faulted depression.

Immediately underlying the lava pond is a sequence of massive flows, pillow lavas, and sheet flows (Units 1256D-2 through 15; 350.3–533.9 mbsf) grouped together as the “inflated flows.” Although rocks exhibiting a number of eruptive styles are included here, the critical criterion for this subdivision is the occurrence of subvertical elongate fractures filled with quenched glass and hyaloclastite (e.g., Sections 206-1256D-21R-1 and 40R-1) at the top of the lava flows. These features indicate flow-lobe inflation, which requires eruption onto a subhorizontal surface with less than a few degrees slope (Umino et al., 2000, 2002); therefore, it is unlikely that such lavas formed on the ridge slope. Such inflation features are not observed in cores deeper in Hole 1256D. Taken together, the cumulative thickness of the lava pond and the inflated flows is ~280 m, close to our preferred estimate of ~300 m for the thickness of off-axis lava flows.

Seismic data from site survey Cruise EW9903 offer significant clues about the expected downhole lithologic variations. The velocity-depth function inferred from seismic refraction (Fig. F7) (A. Harding, unpubl. data) shows a uniform gradient from ~4.8 km/s at the sediment/​basement interface to 5.3 km/s at ~600 msb, with the gradient then sharply increasing and velocity reaching 5.9 km/s at ~800 msb. The gradient abruptly returns to a moderate value, with velocity of 6.5 km/s at ~1250 msb. The gradient decreases gradually, with nearly uniform velocity of ~6.8 km/s between 2000 and 3000 msb. MCS reflection data (Fig. F6) (Hallenborg et al., 2003) show several nearly horizontal reflections with kilometers of horizontal extent to nearly 5.5 s traveltime, or ~800–900 msb. For the upper ~800 m, the relatively lower velocities and horizontal reflection character suggest that flows constitute a substantial fraction of the uppermost crust at Site 1256. At greater depth, the decrease in velocity gradient below ~1250 msb marks the seismic Layer 2/3 boundary. Unfortunately, the gradual nature of the change in gradient means that the depth of this transition cannot be assigned a depth more precisely than 1250–1500 msb. Whether this boundary corresponds the presence of gabbro remains to be tested.

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