Hole 1256A is a single mudline core that recovered the uppermost 2.37 m of the sedimentary section. This interval was resampled with a second mudline core in Hole 1256B, and coring continued in sediment in this hole down to a depth of 250.70 mbsf, where basement was contacted, with roughly 1 m more penetration in and ~5 cm of recovery of the uppermost basaltic basement. The lower portion of the sedimentary section was resampled in the rotary-cored Hole 1256C from a depth of 220.10 mbsf down to 245 mbsf. The combination of these three holes provides a nearly complete sampling of the entire sedimentary sequence above basement at Site 1256. The sediments are clay rich in the upper 40.6 m and become increasingly calcareous with depth, with calcareous nannofossils being the dominant component below 40.6 mbsf, except for some minor more siliceous and diatom-rich intervals (Fig. F17).
Visual observations on the recovered core have been integrated with magnetic susceptibility, density, and color measurements, which are routinely determined during core processing. These data were combined with normal sedimentological criteria (smear slide analysis, sedimentary structures, color, bioturbation, and general appearance) and X-ray diffraction (XRD) results to divide the sediment column into lithologic units. A composite graphical log of the complete sedimentary sequence at Site 1256 (Fig. F18) shows the relationships between sediments and the continuously measured parameters.
The primary density contrast for biogenic oozes is between biogenic calcite and biogenic silica. The density (
) of calcareous nannofossil-rich sediment (e.g., = 1.4-1.6 g/cm3 in Cores 206-1256B-14H through 15H) is greater than that of sediments with a high biogenic silica content (e.g., = 1.1-1.3 g/cm3 in Core 206-1256B-10H). Thus, these continuously recorded gamma ray attenuation (GRA) bulk density measurements provided a first-order estimate of carbonate content, which supplemented our direct observations from smear slides and chemical analyses (see "Inorganic Geochemistry").
The sediments from Site 1256 are divided into two principal lithologic units (Fig. F17; Table T5). Unit I is a clay-rich unit with a few carbonate-rich intervals, and Unit II is predominantly biogenic carbonate. The unit boundaries are based primarily on the relative clastic and biogenic component concentrations, as characterized through smear slide analysis, visual inspection of the core, and changes in color reflectance and physical properties. The subunit boundary within Unit I was based on biostratigraphic and magnetostratigraphic constraints, particularly the location of the Pliocene/Pleistocene boundary, and changes in physical properties.
The clay-rich sediments of lithologic Unit I are divided into two subunits. Subunit IA consists of Pleistocene dark brown to yellow-brown (7.5YR 3/2 to 10YR 5/4) silty clays. The dominant biogenic components are calcareous nannofossils, which compose 10%-25% of the sediment (Fig F19). Subunit IB comprises Pliocene to late Miocene sandy clays to silts with calcareous nannofossil-rich intervals.
Subunit IA comprises ~17.5 m of nannofossil silty clay and clayey nannofossil ooze. The proportions of clay, silt, and sand are variable, and there is a general decrease from 60% nonbiogenic material at the top of the sequence to 40% at the base of the subunit (see "Site 1256 Smear Slides"). Few radiolarians are found in the subunit, but diatoms generally form a minor component (up to 10%). Trace volcanic glass is also observed through the subunit. XRD analyses indicate that the clay-rich intervals are dominated by smectite (Table T6).
A volcanic ash layer consisting predominantly of colorless volcanic glass with common quartz and feldspar is present in interval 206-1256B-3H-2, 34-36 cm (Fig. F20). This ash layer may possibly be correlated with a similar ash layer identified at Site 844 in interval 138-844B-1H-1, 120-126 cm (Mayer, Pisias, Janecek, et al., 1992).
Dominant colors are shades of yellow brown and olive gray that are mottled throughout (10YR 5/4 to 10YR 5/3 and 5Y 5/2). The uppermost 1.5 m of the subunit is very dark brown to dark mottled brown (7.5YR 2.5/2 to 7.5YR 3/2), which is considerably darker than the rest of the sedimentary sequence. Subunit IA has the lowest reflectance of all of the sediments at Site 1256.
Bioturbation is common to abundant throughout. The most common trace fossils are Planolites with tentative identification of Chondrites and Skolithos. Calcified burrows are present in Sections 206-1256B-2H-2 at 100 cm, 2H-7 at 10 cm, and 3H-1 at 55 cm.
The base of Subunit IA is taken as the Pleistocene/Pliocene boundary (see "Hole 1256B" in "Calcareous Nannofossils" in "Biostratigraphy"). At this subunit boundary, the observed change in the relative concentrations of clastic and biogenic components is also apparent in the color reflectance and magnetic susceptibility data.
Subunit IB consists of clayey nannofossil ooze, sandy silty clay, and sandy silty nannofossil ooze. The subunit is characterized by a coarser clastic component than Subunit IA. Diatoms are a minor component of the sediment. As in Subunit IA, radiolarians are sparse, although they are more abundant at the base of Subunit IB.
Dominant colors vary with lithology. The nannofossil oozes are typically light olive gray grading to pale brown (5Y 6/2 to 10YR 6/3) at the base of the subunit, with bioturbation resulting in significant mottling. Intervals dominated by clastic components are darker, in particular Sections 206-1256B-4H-2 through 4H-7. The clastic-rich sediment is olive green to yellowish brown (5Y 5/2 to 10YR 5/4). Localized color banding (olive; 5Y 4/3) is seen in Sections 206-1256B-4H-2 and 4H-4 through 4H-CC. The color reflectance in Subunit IB increases gradually down through the sequence, with a sharp step increase marking the subunit base.
Bioturbation is moderate to abundant throughout. Solid burrows, Planolites, Skolithos, and Zoophycos (Fig. F21) are all common. The top of Sections 206-1256B-5H-1 and 5H-4 contain good examples of Zoophycos cutting other solid burrows. A large (4 cm diameter), continuous calcified burrow intersects the split surface of Section 206-1256B-4H-1 at 15, 35, and 43 cm.
The magnetic susceptibility of Subunit IB is similar in magnitude to that of Subunit IA, but with much less variability. A decrease in the abundance of clay-hosted magnetite grains associated with an increase in nannofossils marks the Unit I/II boundary, and this is reflected in a sharp decrease in chromaticity (a* and b*) (see Fig. F22) at 40.6 mbsf and a decrease (from 30 to 10 in raw SI values) in magnetic susceptibility in the interval 40-45 mbsf.
Lithologic Unit II comprises calcareous nannofossil ooze with varying amounts of clay and other microfossil groups. Diatoms are a significant minor component. Foraminifers and radiolarians are rare throughout the unit. Volcanic glass is absent, although fragmented pumice clasts were identified in the smear slide taken from Section 206-1256B-8H-2 at 67 cm.
The uppermost 35 m of the unit consists predominantly of diatom-nannofossil ooze, with the diatom component varying between 10% and 25%. These biogenic oozes are much paler than the overlying Subunit IB and are dominantly light greenish gray alternating with dark greenish gray to dark gray (5GY 7/1 and 10Y 4/1 to 7.5YR 4/0), typically on a meter-length scale. In some sections there is localized color banding in varying shades of gray.
The diatom content of the sediment increases significantly between 43 and 85 mbsf, resulting in a siliceous biogenic ooze at 85 mbsf. This clayey nannofossil diatom ooze is slightly darker (greenish gray to dark greenish gray; 10Y 7/2 to 10Y 4/1) than the surrounding nannofossil ooze and is lower in density.
With the exception of two intervals, the remainder of the sedimentary section is dominated by almost pure nannofossil ooze (up to 97% nannofossils). This ooze is light greenish gray to white (5G 7/1 to 5Y 8/1) with some dark greenish gray (5G 7/1) intervals and rare purple or dark bluish gray (5PB 4/1) to dark greenish gray intervals. Small dark bluish black blebs, which were observed throughout Core 206-1256B-21X (174.8 mbsf) and in all the deeper cores, are interpreted to be iron sulfide, probably pyrite.
Between 111 and 115 mbsf (Sections 206-1256B-12H-6 through 13H-5) is an unusual diatomite with nannofossils. It consists of a ~4-m-thick diatom mat with a peculiar foliated texture resembling a pile of soggy telephone books (Fig. F23). The XRD analysis for this sediment confirms it to be dominantly opaline sediment. The interval contains very abundant tubular diatom tests, as observed in the smear slide. The diatomite is predominantly light greenish gray to green gray (10Y 8/1 to 10GY 5/1). It has a lower reflectance than the nannofossil ooze and is marked by a significant step decrease in density and an increase in seismic velocity, a reflection of the open cage-like structure of diatoms.
The other anomalous interval spans 140-195 mbsf. In this interval the nannofossil ooze has significant but variable clay and diatom components. There is no obvious change in the visual appearance of the sediment in this interval, but the reflectance and density show greater variability.
Bioturbation is common throughout Unit II. Trace fossils include both solid and rind burrows, including Planolites and Skolithos. The majority of cores contain calcified burrows, generally at the top of Section 1, and as such their location is considered an artifact of the drilling process.
Many of the cores also contain chert nodules, again commonly at the top of Section 1. These chert nodules are generally dark greenish gray (10Y 3/1). Chert was first observed in interval 206-1256B-13H-2, 0-16 cm, at 111.40 mbsf. Distinct chert layers are identified at 111 and 158 mbsf in the wireline logs of Hole 1256C and are characterized by increased resistivity and low porosity (see "Results" in "Downhole Logging" in "Downhole Measurements"). Only 20 cm of sediment was recovered in Core 206-1256B-27X, and it consisted of broken bits of chert and chert nodules with a nannofossil ooze coating, indicating that a chert layer is present somewhere between 230 and 240 mbsf. Recovery of the same interval in Hole 1256C was also extremely poor. The chert layer at this level unfortunately could not be confirmed by downhole logging because a constriction in the borehole prevented the deployment of wireline tools to this depth.
Granular glauconitic bands were identified in Sections 206-1256B-9H-3, 9H-7, 25X-4, and 26X-3. These bands are typically up to 5 mm wide and are dark greenish gray with angular mineral grains up to 2 mm in diameter. In Section 206-1256B-26X-3 glauconite bands are offset by a small fault, possibly induced during coring (Fig. F24).
The nannofossil ooze at the base of the sedimentary section has a turquoise hue, reminiscent of the covers of DSDP Initial Reports volumes. The X-ray diffractogram of this sediment showed it to be dominantly composed of mica, possibly celadonite (Table T6). It is speculated that this may relate to the presence of metalliferous sediment, with celadonite formed by the recrystallization of biogenic opal and hydrothermally derived iron oxides.
In Core 206-1256C-4R, the first core in which igneous basement was recovered, two pieces of red-brown chert were cored. The cherts are silicified recrystallized siliceous sediments with iron oxides, and they may be recrystallized metalliferous sediments.
The color reflectance of Unit II is much less variable than that of Subunit IB, due to the much more homogeneous composition of the sediments and the dominance of calcareous nannofossils. Variations in density are attributed to variable relative concentrations of diatoms and nannofossils. The magnetic susceptibility throughout Unit II is low because of the very low concentrations of magnetic minerals in calcareous oozes. Point-susceptibility meter measurements were discontinued below Core 206-1256B-18H (160.1 mbsf) because the readings were clearly at the noise level of the meter, which has lower sensitivity than the whole-core loop meter. Whole-core susceptibility measurements on the multisensor track (MST) were continued down to basement, with readings commonly at the noise level of the meter. The presence of greenish black (10G 2.5/1) blebs throughout Cores 206-1256-21X through 28X, identified as pyrite spots or other iron sulfides, suggests that any magnetite has been converted to pyrite, which would explain the very low susceptibility and magnetic intensity of this lithologic unit (see "Susceptibility" in "Paleomagnetism").
The trace fossils encountered in the sediments are recorded in detail on individual barrel sheets. Bioturbation is common to abundant throughout the majority of the sediments retrieved from Site 1256. The trace fossils typical of Unit I are solid burrows, commonly Planolites with minor Zoophycos and rare Chondrites. The carbonate-rich lithologies of Unit II contain common Planolites with Chondrites and Skolithos. Carbonate concretions up to 2 cm in diameter in Unit II are identified as calcified burrows.
Sediments recovered above basaltic basement at Site 1256 during Leg 206 provide a continuous sedimentary record from the Quaternary through the middle Miocene. Calcareous nannofossils were used for developing the biostratigraphic framework using the zonation schemes of Martini and Müller (1986) and Okada and Bukry (1980). Biostratigraphic assignments were made to core catcher samples and additional samples selected to refine the biostratigraphy. The interval (core and section) and depth (mbsf) constraints for calcareous nannofossil events recognized at Site 1256 are reported in Table T7. The depths of epoch boundaries are given in Table T8, along with the same boundaries from Sites 844 and 845. Nannofossil ranges from all core catcher samples are given in Table T9.
Calcareous nannofossils are generally abundant and moderately to well preserved at Site 1256. Nannofossil assemblages, however, are affected to different degrees by alteration, mainly by etching above Sample 206-1256B-3H-CC, by dissolution and overgrowth below Sample 206-1256B-13H-CC, and by fragmentation throughout the sequence. No significant reworking of nannofossils was apparent, and all nannofossil events are recognized in normal stratigraphic sequence in these sediments. More than a dozen nannofossil datums were determined based on examination of core catcher samples and other selected samples, providing modest biostratigraphic resolution for the Pleistocene through middle Miocene.
Hole 1256A consists of a mudline core only (0-2.17 mbsf) and contains uppermost Quaternary nannofossil assemblages assigned to Zone NN21. Calcareous nannofossils are common throughout the sequence and are moderately preserved. The main assemblage is composed of placoliths (Emiliania huxleyi and Gephyrocapsa spp.) with a highly diverse subordinate assemblage. Reworked specimens, mainly discoasterids of Pliocene to Miocene age, were rarely encountered. E. huxleyi, the first occurrence (FO) of which defines the base of Zone NN21 (0.26 Ma), was observed throughout the sequence upward from the lowermost Sample 206-1256A-1H-CC.
Core was recovered in Hole 1256B from the mudline to 250.70 mbsf. The sediments recovered were assigned to Zone NN21 of the Pleistocene through Zone NN5 of the middle Miocene. Cores 206-1256B-1H through 4H (0-34.86 mbsf) yield mostly moderately preserved nannofossil assemblages. Nannofossils are common throughout the Pleistocene and Pliocene sequence, but several intervals in Cores 206-1256B-3H and 4H are barren of nannofossils (Samples 206-1256B-3H-2, 35 cm, and 40-42 cm; 3H-6, 40-42 cm; 4H-1, 40-42 cm; and 4H-5, 40-42 cm). Therefore, the proper position of some biostratigraphic events defining zonal boundaries may be within the barren intervals and can only be resolved by future biostratigraphic analysis of other fossil types. There is a gradual increase in nannofossil abundance from Sample 206-1256B-5H-CC (44.31 mbsf), with a significant increase in abundance and an improvement in preservation from Sample 206-1256B-8H-CC (72.80 mbsf). The preservation deteriorates with increasing depth from Sample 206-1256B-14H-CC (129.82 mbsf) to the base of the sediment section. Nannofossil assemblages here are affected by strong overgrowth, which mainly affected discoasterids and sphenoliths. This, however, did not inhibit the identification of the middle Miocene nannofossil zonal boundaries, as the zonal markers could still be recognized even in specimens suffering from overgrowth.
The Pleistocene includes the interval from 0 to 15.75 mbsf. Moderately preserved Pleistocene nannofossils were recovered from Samples 206-1256B-1H-2, 67 cm, through 2H-CC. The Pleistocene assemblages older than 0.26 Ma are characterized by the dominance of gephyrocapsids. Reworked earlier Neogene taxa such as discoasterids and Sphenolithus spp. were found in low numbers throughout the Pleistocene samples.
The base of Zone NN21, marked by the FO of E. huxleyi (0.26 Ma), was recorded between Samples 206-1256B-1H-2, 67 cm, and 1H-4, 60 cm. The last occurrence (LO) of Pseudoemiliania lacunosa (0.46 Ma) was recorded between Samples 206-1256B-1H-CC and 2H-2, 113 cm, bounding the bottom of Zone NN20 and the top of Zone NN19. Further quantitative analysis of relative abundance of E. huxleyi and the different Gephyrocapsa morphotypes may provide a higher biostratigraphic resolution for the Pleistocene.
Samples 206-1256B-3H-4, 40-42 cm (20.50 mbsf), through 5H-2, 40-42 cm (36.50 mbsf), were assigned to the Pliocene. The Pliocene sediments contain moderately preserved nannofossils in most samples. The abundance of nannofossils is generally low relative to deeper intervals, and Samples 206-1256B-3H-2, 35 cm, and 40-42 cm; 3H-6, 40-42 cm; and 4H-5, 40-42 cm, are barren. The Pliocene nannofossil assemblages are dominated by different morphotypes of reticulofenestrids, sphenoliths, and discoasterids. Other taxa are mainly Coccolithus spp. and Calcidiscus spp.
The LO of Discoaster brouweri (1.96 Ma) was used to determine the top of Zone NN18 between Samples 206-1256B-2H-CC and 3H-4, 40-42 cm. After several nannofossil-barren intervals (Samples 206-1256B-3H-2, 35 cm, and 40-42 cm, and 3H-6, 40-42 cm), the LO of Discoaster pentaradiatus (2.52 Ma) was recorded between Samples 206-1256B-3H-5, 83 cm, and 3H-CC, assigning the level to Zone NN17. The top of Zone NN16 is marked by the LO of Discoaster surculus (2.53 Ma) between Samples 206-1256B-3H-CC and 4H-2, 40-42 cm. The LO of Reticulofenestra pseudoumbilica (>7 µm) was used to define the Zone NN16/NN15 boundary (3.82 Ma) between Samples 206-1256B-4H-4, 40-42 cm, and 4H-6, 40-42 cm. The LO of Sphenolithus abies/neoabies (3.66 Ma) was found in Zone NN16 between Samples 206-1256B-4H-2, 40-42 cm, and 4H-4, 40-42 cm. Amauroliths are very rare throughout the sequence. Among the 61 samples examined from Hole 1256B, only one specimen of Amaurolithus primus was observed in Sample 206-1256B-5H-2, 40-42 cm. Therefore, the top of Zone NN14/NN13 and the boundary between Subzones NN11b and NN11a, which is defined as the LO of Amaurolithus spp. and the FO of A. primus, respectively, remain undetermined. The LO of Discoaster quinqueramus (5.6 Ma) at the top of Zone NN11, which approximates the Pliocene/Miocene boundary (5.32 Ma), is between Samples 206-1256B-5H-5, 40-42 cm, and 5H-6, 93-100 cm.
Sediments recovered from Sample 206-1256B-5H-5, 40-42 cm (41.00 mbsf), through the bottom of Hole 1256B (250.70 mbsf) contain Miocene nannofossils. The generally well- to moderately preserved Miocene nannofossil assemblage consists mainly of placoliths, diverse discoasterids, and sphenoliths.
The FO of Discoaster berggrenii was used to define the boundary between Zones NN10 and NN11 (8.6 Ma) between Samples 206-1256B-9H-2, 80 cm, and 9H-CC. The range of Discoaster hamatus defines the boundaries of Zone NN9 (9.4-10.38 Ma). Its top occurrence was recorded between Samples 206-1256B-9H-CC and 10H-4, 100 cm, and the bottom occurrence between Samples 11H-CC and 12H-4, 60 cm. The basis of Zone NN8 is marked by the FO of Catinaster coalitus (10.9 Ma) between Samples 206-1256B-12H-4, 60 cm, and 12H-5, 52 cm. This event approximates the upper/middle Miocene boundary (11.2 Ma). The FO of Discoaster kugleri, marking the boundary between Zones NN7 and NN6 (11.8 Ma), was recorded between Samples 206-1256B-16H-CC and 17H-3, 112 cm. D. kugleri can still be distinguished from Discoaster sanmiguelensis and other discoasterids in overgrown specimens by the diagnostic six short rays and broad, flat central area. The LO of Cyclicargolithus floridanus (13.2 Ma) between Samples 206-1256B-23X-3, 97 cm, and 23X-CC provides a further datum to subdivide Zone NN6, whereas its presence in Sample 206-1256B-20X-CC is considered as reworked. The LO of Sphenolithus heteromorphus (13.6 Ma), altered by overgrowth but recognizable by special diagnostic characteristics, was recorded as the lowermost biostratigraphic zonal event in the sequence, defining the top of Zone NN5 between Samples 206-1256B-24X-CC and 25X-2, 90 cm. Helicosphaera ampliaperta, the LO of which marks the top of Zone NN4 (15.6 Ma), was absent from all samples including the lowermost Sample 206-1256B-28X-CC. Assuming a constant sedimentation rate of 36 m/m.y., based on datums from below 100 mbsf, implies a basement age of ~14.6 Ma (see Fig. F25), consistent with assignment to Zone NN5 of the middle Miocene. This age is reasonably consistent with the age of 15.1 Ma inferred from magnetic anomalies (see "Regional Geology and Tectonic Setting" in the "Leg 206 Summary" chapter).
The geological and tectonic background at Site 1256 seems similar to nearby Sites 844 and 845 in the Guatemala Basin as indicated by close correlation of the epoch boundaries among these sites (Table T8). Calculated sedimentation rates from the age-depth plot are high in the middle Miocene (36 m/m.y.), decrease drastically in the upper Miocene (8-14 m/m.y.), and are low in the Pliocene and lower Pleistocene (6 m/m.y.), but increase afterward (12 m/m.y.) (see Fig. F25). This pattern of sedimentation rate variations was also observed at Sites 844 and 845, with the event that occurred at the end of the middle Miocene being referred to as a carbonate crash (Farrell et al., 1995). The high sedimentation rate in middle Miocene can be attributed to very high productivity while the site was near the paleoequator and to complete preservation on young, shallow seafloor. In general, the more recent slower rates can be attributed to lower productivity away from the equator and to partial dissolution after the seafloor subsided. In detail, the late Miocene pattern is much more complex. Farrell et al. (1995) interpreted that the late Miocene sedimentation rates in the tropical eastern Pacific were controlled by large variations in the CCD, with nearly complete carbonate dissolution during brief intervals when the CCD shoaled by up to 1400 m. Further shore-based work with higher-resolution sampling will allow us to test whether Site 1256 shows synchronous variations with other nearby sites.
We use progressive alternating-field (AF) demagnetization of the sedimentary split-core sections and discrete samples along with rock magnetic experiments to characterize the paleomagnetic signal and resolve the magnetization components recorded in the recovered core. The component interpreted to record the magnetization at or near the time of deposition is then used to construct a magnetostratigraphy of the sedimentary overburden at Site 1256. Whole-core and split-core susceptibility measurements and anhysteretic remanent magnetization (ARM) and isothermal remanent magnetization (IRM) experiments provide additional information for characterizing the rock magnetic properties of the sediments.
All split-core and discrete samples have a sizable drilling overprint, which is characterized by a steep downward direction, but also by a radial-horizontal component that points toward the center of the core. In the ODP core orientation system, the latter results in a strong bias in the declinations toward 0° for archive-half measurements. Initial natural remanent magnetization (NRM) measurements are thus characterized by inclinations of more than +60° and declinations of ~0°. In general, 15- to 20-mT AF demagnetization removes most or all of the drilling overprint but also reduces the magnetization by ~80%. As has been shown in prior studies (e.g., Shipboard Scientific Party, 2002), the drilling overprint can be more demagnetization resistant along the periphery of the core, sometimes requiring AF demagnetization of up to 60 mT to fully remove the overprint. Small biases may thus persist in some of the split-core measurements, for which AF demagnetization did not exceed 40 mT.
Biases or completely bogus data are also associated with remanence measurements over intervals disturbed or deformed by coring. Similarly, magnetic edge effects, which can be large when measurements are within ~5 cm of the edge of a section or edge of a void, can give biased results. To avoid interpreting results in these regions, we recorded where coring disturbance was severe enough to be visibly identified and where voids or gaps were located in the APC cores (Table T10). Paleomagnetic and susceptibility data from these intervals were removed prior to interpretation.
Following removal of most or all of the drilling overprint and excluding edge effects and intervals highly disturbed by drilling, the sediments above 110 mbsf (Cores 206-1256B-1H through 12H) have stable remanent magnetizations that provide a record of the geomagnetic field at or near the time of deposition. Below this, stable directions are difficult or impossible to resolve. The main factors appear to be weak magnetic intensities, with associated high variability in inclinations and declinations, and deformation related to coring (mainly XCB coring). The intensity decrease downhole coincides with a downhole increase in carbonate content. The intensity decrease may also be related to reduction diagenesis, as the major drop in intensity is at least qualitatively associated with more common iron sulfides, apparent by the presence of black patches on the split-core surfaces below 110 mbsf. The direct impact of deformation caused by XCB coring is less obvious because the magnetic signal is already dominated by noise before XCB coring was initiated. The shallow inclination expected for the sediments, which were deposited near the paleoequator, along with the lack of azimuthal orientation for XCB cores and the relative azimuthal rotation of core that occurs on the centimeter to decimeter scale as XCB cores are recovered, means that little useful paleomagnetic information could be extracted from the XCB cores even if a stable direction had been present.
Relative changes in whole-core susceptibility, measured every 2.5 cm by a loop susceptibility meter on the MST, and the archive-half core susceptibility, measured every 2.5 cm by a point susceptibility meter on the archive multisensor track (AMST), are in good agreement (Figs. F26, F27). The raw meter values from the MST loop meter are ~2.4 times higher than those measured by the AMST point meter. Thus, to correct the AMST raw values to true volume susceptibilities in SI units, the values would be multiplied by ~2.4 and then by a correction factor similar to that applied to the raw MST data, which are multiplied by ~0.7 x 10-5 to convert to SI volume units (see "Susceptibility" in "Instruments and Measurements" in "Paleomagnetism" in the "Explanatory Notes" chapter). In the following, we give susceptibility values in raw meter units.
The sensitivity of the point susceptibility meter, which is again about a factor of two lower than that of the loop meter, is at or very near the noise limit of the instrument for the sediments below ~90 mbsf (Fig. F26). Raw meter readings were commonly negative below this depth, which may be caused by the diamagnetism of the carbonate-rich sediments but is probably also partly an artifact of this device (the true zero for the meter could be offset by a few meter values). In any case, we did not make additional point susceptibility measurements below 160 mbsf after it became clear that the signal was at the noise level of the meter. The loop meter, with its higher sensitivity, provides clear evidence that the susceptibility was very low from 110 to 205 mbsf, with raw meter values averaging -0.1 over the interval.
Whole-core susceptibility values are highest in the upper 40.6 mbsf (above the section break between Sections 206-1256B-5H-4 and 5H-5), averaging 28.7 raw meter units (Fig. F26). Within this interval, the susceptibility has more variability in the upper ~20 mbsf. This corresponds to the lithologic Subunit IA/IB boundary at 17.48 mbsf. The susceptibility decreases gradually downhole from 40.6 to ~44 mbsf, where the values then average 6.0 raw meter units down to 110 mbsf, with values decreasing slightly over the interval from 44 to 110 mbsf. Susceptibilities average -0.1 from 110 to 205 mbsf and then increase to an average of 7.5 raw meter units from 205 mbsf to basement.
The 40.6-mbsf boundary is notable as the contact between more clay-rich sediments above and more calcareous-rich sediments below. The contact defines the boundary between lithologic Units I and II (see "Description of Units" in "Lithostratigraphy"). The abrupt downhole decrease in a* and b* chromaticity, which occurs at this contact and was used to define the unit boundary, indicates a downhole decrease in the amount of red and yellow, respectively, and an increase in green and blue, respectively (Fig. F22). As described in the visual core descriptions, the color changes downhole from pale brown and olive to greenish gray. The magnetic intensity also increases uphole by more than a factor of five across this boundary (Fig. F22). The higher susceptibility and magnetic intensity above 40.6 mbsf most likely result from a higher abundance of titanomagnetite in the terrigenous sedimentary component (clays) relative to that in the pelagic component (calcareous nannofossils, diatoms, and radiolarians).
The downhole drop in susceptibility at 110 mbsf corresponds to a downhole increase in carbonate content over an interval from ~104 to 117 mbsf, with the average carbonate content being ~30 wt% above and ~80 wt% below this interval (Fig. F22). The downhole increase in susceptibility at 205 mbsf is associated with more common glauconite and chert downhole, although it is not clear that either are the direct cause of the susceptibility increase.
During ODP cruises where multiple holes are cored at a site with overlapping cored intervals, a composite or common depth scale is constructed by correlating features or physical properties between holes (e.g., Acton et al., 2001). The resulting stratigraphic framework allows observations, measurements, and samples taken from one hole to be directly related to other holes cored at the site. The magnetic susceptibility data are commonly the data that are most easily correlated between holes when creating a composite sedimentary section.
For Site 1256, only the top couple of meters below the mudline and a few tens of meters above basement were multiply cored. Recovery was very low for the RCB cores from above basement in Hole 1256C. Therefore, no attempt is made to correlate them with the XCB cores from Hole 1256B. Correlation between the mudline cores will, however, facilitate their future use. Comparison of the susceptibility records from Core 206-1256A-1H (0-2.37 mbsf), which is the only core collected from Hole 1256A, to Core 206-1256B-1H (0-6.14 mbsf) indicates that the two mbsf depth scales differ by no more than about the distance between one or two susceptibility measurements (Fig. F27). This equates to the two scales differing by ~2.5 to 5 cm, with distinctive correlative susceptibility anomalies being consistently a few centimeters deeper in Hole 1256A than in Hole 1256B. The difference is so small that it could be derived from any number of minor distance measurement errors, from the initial curation length measurement to how the sections were positioned in the MST. Thus, the mbsf depth scale from Hole 1256A is nearly equivalent to that of Hole 1256B and no new meters composite depth scale need be constructed.
NRM was routinely measured every 2.5 cm along all archive-half sections from the sediments from Hole 1256A and on Hole 1256B sediments through Section 206-1256B-17H-3 (153.10 mbsf) before and after demagnetization up to 30 or 40 mT. Because the quality of the data degraded downhole and core flow was being slowed by the pace of measurements in the paleomagnetism laboratory, we expanded the measurement interval to 5 cm and continued with AF demagnetization up to 40 mT for sediments from 153.10 to 188.50 mbsf (Sections 206-1256B-17H-4 through 22X-3). When it became obvious that the AF demagnetization results further downhole would not be useful for establishing a magnetostratigraphy, we ceased demagnetization treatments. Thus, for all sediment archive-half sections from below 188.5 mbsf, only NRM was measured.
After removing data collected in disturbed intervals and near voids, the clean data set (Tables T11, T12, T13) is characterized by several distinct changes in magnetic properties, similar to those described above for susceptibility. Because the cores are fairly uniformly remagnetized by the drilling overprint, somewhat similar to a viscous low- to medium-field IRM, both the NRM intensity and susceptibility are primarily measures of magnetic mineral concentration. Thus, the concentration of magnetic minerals is highest in the upper 40.6 mbsf, intermediate between 40.6 and 110 mbsf, lowest between 110 and ~160 mbsf, and intermediate to low and highly variable from ~160 mbsf to basement (Fig. F27).
The combination of (1) weak intensities; (2) drilling overprints; (3) drilling deformation and the lack of azimuthal orientation related to XCB coring; and (4) possibly some level of reduction diagenesis, which may have destroyed the primary magnetization signal, results in a paleomagnetic signal that is dominated by noise below 110 mbsf (Fig. F28). The data below 110 mbsf are therefore not used in magnetostratigraphy or paleomagnetic interpretations that follow.
Paleomagnetic directions from sediments above 110 mbsf before demagnetization are dominated by a steep downward-directed inclination caused by the drilling overprint (Fig. F28). Declinations before demagnetization and before azimuthal reorientation are biased toward 0°. This is related to the radial drilling overprint. Even with this bias, the declinations show signs that a remanent component exists, as evidenced in the way that they cluster for each core at values different from other cores and different from 0° (Fig. F28).
Demagnetization of the split cores up to 30 or 40 mT removed most but not all of the overprint. Much of the steep drilling overprint and ~80% of the total NRM magnetization were removed by 15 to 20 mT, typically resulting in directions with shallow, but dominantly positive, inclinations (Fig. F29). This shallow component, although still partially obscured by the drilling overprint, is interpreted as the characteristic remanent magnetization (ChRM) acquired by the sediments during or shortly after deposition.
We attempted to refine the estimates of the ChRM using principal component analysis (PCA) and stable endpoint averages as described in "Data Reduction and Analysis" in "Paleomagnetism" in the "Explanatory Notes" chapter (Table T14). Because the overprint was not fully removed by demagnetization at 30 to 40 mT, the highest demagnetization steps used on split-core sections, neither method is totally satisfactory for estimating the ChRM direction. Biases in the directions of several degrees to tens of degrees occur, depending on how well the drilling overprint is removed, which varies over the core. Angular distances between the two estimates are <15° for more than half of the data (Fig. F30). The declinations estimated by the two methods are similar enough that little ambiguity exists in interpreting the magnetic polarity, although slight variations in the locations of reversal boundaries may occur, as can be seen for the Jaramillo Subchron in Figure F31.
Typically, one sample was collected from each sediment core from Hole 1256B and subjected to progressive AF demagnetization in 2-mT steps from 0 to 20 mT and then 5-mT steps from 20 to 80 mT (Table T15). Samples were taken from the center portion of the working-half cores, where drilling disturbance and overprinting are reduced relative to sediment near the core periphery. That and the higher demagnetization allow the ChRM to be much better resolved in the discrete samples (Figs. F32, F33, F34, F35). The more detailed demagnetization also illustrates that the drilling overprint dominates the NRM signal but is possibly completely removed by 15-20 mT in the discrete samples.
The inclinations from normal polarity zones (e.g., Figs. F32, F35) are generally steeper than those from reversed polarity zones (e.g., Figs. F33, F34), which might indicate that the steep positive drilling overprint is not completely removed even at high AF demagnetization or, perhaps as likely, that a recent Brunhes overprint is not completely removed. Alternatively, real differences may exist between the polarity states. The vector demagnetization diagrams also illustrate that the linear demagnetization paths do not decay directly to the origin (Fig. F34). Such offsets have been observed during previous cruises (Shipboard Scientific Party, 2003) and likely result from a very weak ARM imparted to the samples by the in-line AF demagnetizer that becomes more pronounced at higher demagnetization fields. Because of this, the stable endpoints will not give accurate estimates of the ChRM for the discrete samples. The PCA results are thus preferred for the discrete samples (Table T16). As noted for the split-core samples, the ChRM cannot be resolved below 110 mbsf (e.g., Fig. F36).
Circumstantial evidence about the magnetic mineral that carries the remanent magnetic signal can be obtained from the susceptibility and remanence measurements and from the ARM and IRM experiments. Titanomagnetites (including magnetite) are by far the most common carrier in marine sediments, and it would be surprising if they were not present. AF demagnetization results indicate the coercivity of the magnetic carrier is consistent with it being titanomagnetite, as well as perhaps maghemite or pyrrhotite. Curie temperature analyses and low-temperature magnetic experiments to be conducted postcruise will better determine the dominant magnetic mineral.
Assuming that titanomagnetite is the dominant magnetic mineral, we conducted Lowrie-Fuller tests (Lowrie and Fuller, 1971). The test is useful for discerning magnetic grain size of titanomagnetites through comparison of the demagnetization of ARM to that of a saturation IRM. In this test, the ARM is intended to represent a magnetization similar to a low-field thermal remanent magnetization (TRM) and the saturation IRM is intended to represent a magnetization similar to a strong-field TRM (see pp. 306-309 of Dunlop and Özdemir, 1997). If these conditions are met, the size of the alternating field required to remove the ARM will increase as the grain size decreases, whereas that required to remove the IRM will decrease. Commonly, the median destructive field (MDF), which is the alternating field required to remove one-half of the magnetization, is used for comparison. When the ratio of the MDF for the ARM to that for the IRM is ~1 or higher, the grain size is less than ~10 µm, or in the single-domain and/or pseudosingle-domain grain size for crushed natural magnetites. As the ratio decreases, the grain size progresses to larger pseudosingle-domain size to multidomain size.
Thus, after completion of the NRM measurements, we imparted an ARM (50-µT biasing field and 100-mT AF demagnetization) to each discrete sediment sample and then measured the remanence of each after 0- to 80-mT demagnetization in 5-mT steps (Table T17). Following this, we imparted an IRM (1 T) to each of the samples and again measured the remanence of each after 0- to 80-mT demagnetization in 5-mT steps (Table T18). Comparison of the MDF of the ARM and IRM indicates that the magnetic grain size is smaller in the upper clay-rich interval (above 40.6 mbsf) where the ratio of the MDF is ~1 (Fig. F37). Below this and down to 110 mbsf, the Lowrie-Fuller test would indicate that the grain size falls into the pseudosingle-domain and multidomain grain sizes. Below 110 mbsf, the test is ambiguous because the ARM imparts little magnetization, with subsequent demagnetization being dominated by noise (see Sample 206-1256B-13H-3, 62 cm, in Fig. F37).
The azimuth of each APC core was reoriented to geographical north using a combination of Tensor tool orientation information (see "Core Orientation" in "Paleomagnetism" in the "Explanatory Notes" chapter) or, for Cores 206-1256A-1H and 206-1256B-1H, 2H, and 9H, by assuming that the mean Brunhes declination was 0°. The angles that need to be added to the observed declination to correct them to geographic north are given in Table T19.
No Tensor tool data were collected for the only core from Hole 1256A or the upper two cores from Hole 1256B because the BHA was not sufficiently buried in the unconsolidated sediments near the seafloor to allow for accurate measurement. The Tensor orientation was in error for only Core 206-1256B-9H, as was evidenced by declinations that were preferentially aligned along 80° and 260° after reorienting the core. Based on magnetostratigraphic interpretation of the adjacent cores, an arbitrary 100° would need to be added to the Tensor tool correction to align the declinations preferentially along 0° and 180°. In other intervals, the Tensor tool orientation appears to be off by as much as ±30°, although part of the variation may be geomagnetic. In any case, the orientation information is good enough to allow the declinations to be used for magnetostratigraphic interpretations, which was our main goal in orienting the sediment cores.
Because the inclination is near zero at Site 1256 and because the drilling overprint biases the inclination toward positive values regardless of whether the sediment has reversed or normal polarity, determining the magnetic polarity from the inclinations is difficult at best. The declination of the ChRM after core orientation gives a clear indication of polarity, where values near 0° are indicative of normal polarity and those near 180° are indicative of reversed polarity. In addition, polarity can often be inferred with confidence by comparing the depth intervals between ~180° changes in declinations with the time intervals between known reversal ages.
The magnetostratigraphic record from Site 1256 appears to record nearly all chrons and subchrons from Chron C1n (Brunhes Chron; 0.0-0.780 Ma) through most of Subchron C5n.2n (9.920-10.949 Ma) (Figs. F38, F39; Table T20). For depths of 0-40 mbsf, the noise level is low and Chron C1n (0-0.780 Ma) to the termination of Chron C3r (5.230 Ma) can be identified from the depth ranges of uniform declination, even without using core orientation data (Fig. F39). Deeper intervals show more common noisy intervals with scattered declinations, rendering identifications less confident. Subchrons C3An.1n and C3An.1r, between 49.50 and 50.60 mbsf, are clear, although Chron C3r above this contains significant noise and the subchrons and chrons in the interval from 50.6 to 71.56 mbsf are very poorly resolved (Fig. F39). The next identifiable polarity sequence that can be clearly correlated with the geomagnetic polarity timescale begins at 71.56 mbsf with Subchron C4n.1r and extends down to the termination of Subchron C5n.2n (9.920 Ma) at 92.53 mbsf. The onset of Subchron C5n.2n (10.949 Ma) cannot be identified with confidence, as it is either very near 110 mbsf or below, either in the coring gap between Cores 206-1256B-12H and 13H or lower in the section where the ChRM could not be resolved.
The age assignments from Site 1256 magnetostratigraphy are in very good agreement with the biostratigraphic constraints from calcareous nannofossil datums (see Fig. F28). The magnetostratigraphy provides higher resolution in the late Miocene up to the Brunhes/Matuyama reversal at 0.78 Ma, whereas the biostratigraphy provides constraints within the late Pleistocene, where there is an absence of polarity reversals, and in the middle Miocene, where the paleomagnetic signal was noisy.
Sedimentation rates are estimated from the biostratigraphic datums and the magnetostratigraphic record derived from the sedimentary sequence at Site 1256 (see "Biostratigraphy" and "Paleomagnetism"). Ages derived from these two data sets are in very good agreement (Fig. F25). The magnetostratigraphic ages provide higher resolution in the late Miocene up to the Brunhes/Matuyama reversal at 0.78 Ma, whereas the biostratigraphy provides constraints within the late Pleistocene, where there is an absence of polarity reversals, and in the middle Miocene, where the paleomagnetic signal was noisy.
Sedimentation rates vary from ~6 to 36 m/m.y., with the rate being about four times faster in the middle Miocene than the average rate during the late Miocene to present (Fig. F40). Linear sedimentation rates were estimated for five intervals within which the slope of the age vs. depth data from the magnetostratigraphy and biostratigraphy was constant or nearly so (Fig. F25; Table T21). The linear sedimentation rates are 11.7 m/m.y. from 0 to 1.2 Ma (0-13.87 mbsf), 6.3 m/m.y. from 1.2 to 5.23 Ma (13.87-39.7 mbsf), 13.6 m/m.y. from 5.23 to 8.07 Ma (39.7-73 mbsf), 8.5 m/m.y. from 8.07 to 9.92 Ma (73-95 mbsf), and 36.4 m/m.y. from 9.92 to 13.6 Ma (95-212.65 mbsf). When the 36.4-m/m.y. rate is extrapolated to the basement at 250.7 mbsf, the mean age obtained for the basement is 14.6 Ma, which is consistent with the ~15-Ma age of the oceanic crust estimated from marine magnetic anomalies (see "Regional Geology and Tectonic Setting" in the "Leg 206 Summary" chapter).
A similar pattern of sedimentation rate variation was observed at Sites 844 and 845, with the event that occurs at the end of middle Miocene being referred to as a carbonate crash (Farrell et al., 1995; Lyle et al., 1995). The carbonate crash extended from ~11.2 to 7.5 Ma, with the crash nadir at ~9.5 Ma in the equatorial Pacific (Farrell et al., 1995). Presumably, the beginning of the carbonate crash at Site 1256 corresponds to the base of the diatom mat (115 mbsf), which has an age of ~10.9 Ma. The carbonate content in this interval is ~12 wt%, whereas it averages ~79 wt% below (see "Calcium Carbonate and Organic Carbon" in "Sediment Chemistry" in "Inorganic Geochemistry"). Although the density of carbonate analyses for Site 1256 is relatively low, the crash nadir would presumably occur in the vicinity of the extreme carbonate low (0.25 wt%) at 89.55 mbsf, which has an age of ~9.6 Ma (within Subchron 4Ar.2n). Throughout the late Miocene, the carbonate content is less than one-half what it was below the diatom mat in the middle Miocene, and the sedimentation rates remained low from the beginning of late Miocene to present. The high sedimentation rate in the middle Miocene can be attributed to the very high productivity while the site was near the paleoequator and to complete preservation on young, shallow seafloor. The more recent slower rates can be attributed to lower productivity away from the equator and to partial dissolution after the seafloor subsided. In detail, the late Miocene pattern is much more complex and the sedimentation rates during this period in the tropical eastern Pacific and extending into the Caribbean were controlled or at least influenced by large variations in CCD depth, such as the carbonate crash (Farrell et al., 1995; Lyle et al., 1995; Shipboard Scientific Party, 1997b; Roth et al., 2000).
Variations with depth in the concentrations of dissolved species in sedimentary interstitial waters are dependent on a number of competing processes. To understand these processes, it is necessary to identify the reactions occurring within the sediments including the dissolution and recrystallization of biogenic phases and the growth of authigenic minerals. Also required are the quantification of vertical and lateral diffusion and/or advection of chemical tracers in the fluids. This requires detailed knowledge of the physical properties of the sediment and fluid, fluid velocities, and the composition of basement formation fluids (Mottl, 1989). Detailed modeling of tracer transport by advection or diffusion is beyond the scope of shipboard investigations, but in this section we use pore water profiles to identify reactions occurring within the sediments and discuss the mobility of chemical species in the fluid phase by either diffusion or advection.
Following Mottl (1989), the major reactions that influence sediment pore water profiles are
We collected 104 interstitial water samples from 28 cores at Site 1256. The first 100 were sampled from Hole 1256B, ranging from 2.2 to 226.6 mbsf, and the other 4 were from Hole 1256C from near basement at depths ranging from 240.4 to 245.0 mbsf. We completed shipboard analyses on all of the samples, with analyses from these two holes considered as a single depth profile. As a result of limited argon supply for the analytical laboratory, only 27 of the 104 samples (about 1 per core) were analyzed for elemental composition. Interstitial water samples from 2.2 to 40.6 mbsf were recovered from clays containing varying amounts of calcareous nannofossils, diatoms, and radiolarians that compose lithologic Unit I. The remainder of the interstitial water samples, from 40.6 mbsf to basement at 250.7 mbsf, were extracted from biogenic carbonate/siliceous sediments of Unit II, including the deepest sample (206-1256C-4R-1, 0-5 cm) at 245 mbsf. As a result of poor recovery from Hole 1256C, some samples from this interval were taken from split-core sections of Hole 1256B 1-2 days after they were recovered. As discussed below, interstitial water samples from Core 206-1256B-26X appear to be slightly diluted and could have been contaminated during splitting. Most of the data from this interval are, however, consistent with the general chemical trends observed throughout the sediment column.
Chemical gradients are variable with depth but display distinct trends, which principally reflect diffusion profiles between seawater and basement, overprinted by diagenetic reactions occurring in the sediment column. A 1-m-thick chert bed located at 158 mbsf (see "Results" in "Downhole Logging" in "Downhole Measurements") appears to act as a diffusion barrier, and a sharp step in some chemical concentrations is observed at this depth. Interstitial water profiles at Site 1256 are similar to those at Site 504, occupied during ODP Leg 69 (Mottl et al., 1983). The analyses of interstitial waters collected during Leg 206 are reported in Table T22.
The salinity, sodium, and chloride concentrations (Fig. F41) are relatively uniform downhole, whereas pH decreases slightly, indicating no significant freshwater or brine input. Section 206-1256B-26X-2 (225 mbsf) displays a significantly lower chloride concentration (532 mM) than most samples from this hole, which could result from dilution by cutting fluids during splitting. The Na/Cl ratio is invariant (0.857 ± 0.007) except for two anomalous samples at 9.77 and 187 mbsf (Fig. F41).
High alkalinity and low sulfate concentrations are often indicative of bacterial degradation of organic matter in an organic-rich sediment column, as observed at many DSDP and ODP sites. Commonly, bacterial sulfate reduction utilizes organic matter to produce bicarbonate, the main component of alkalinity. However, at Site 1256, alkalinity gradually decreases downhole, with a net decrease of 51% and higher variability between 50 and 159 mbsf (Sections 206-1256B-6H-4 through 18H-1) (Fig. F42), whereas sulfate concentrations (Fig. F42) decrease steadily from 27 mM at 2.2 mbsf to 21 mM at 245 mbsf, with the lowest concentration of 17.5 mM at 189.9 mbsf (Sample 206-1256B-22X-2, 145-150 cm). Sulfate reduction is occurring, as evidenced by the decreasing sulfate concentrations with depth and formation of secondary pyrite (see "Unit II" in "Description of Units" in "Lithostratigraphy"), but microbial reduction of sulfate is clearly not dominating the alkalinity profile. The observed decrease in alkalinity appears to be a response to calcium carbonate precipitation, whereby calcium combines with bicarbonate to produce calcium carbonate. This is also evident in strontium concentrations (Fig. F43) and is discussed in a later section.
The minimum sulfate values are ~20 mM and are unlike those from nearshore settings where higher organic carbon contents and higher bulk sedimentation rates are common, such as the Gulf of California (Gieskes et al., 1982), the Peruvian margin (Suess, von Huene, et al., 1988), and the West African margin (Murray et al., 1998). In fact, the sediments from Site 1256 are characterized by low (<3 wt%) organic carbon content (see "Calcium Carbonate and Organic Carbon" in "Sediment Chemistry"). The dissolved organic content (Fig. F42) in the top 35 mbsf (through Section 206-1256B-4H-7) steeply decreases to <0.5 mM and remains relatively constant down to basement. The dissolved organic carbon is indicative of organic degradation, and in this case, organic degradation is minimal because of the low abundance of dissolved organic carbon and organic carbon concentrations in sediment. The overall concentration of organic carbon in sediment is <2 mM, and the maximum concentration is 3.5 mM (see "Calcium Carbonate and Organic Carbon" in "Sediment Chemistry"). Low ammonia concentrations are also consistent with low organic degradation (Fig. F42). Phosphate was below the detection limit (<0.1 µM), also consistent with the low organic content.
Sulfate concentrations lower than seawater values are indicative of a reducing environment, which is also suggested by secondary pyrite observed in the sediments. The presence of sulfate in the interstitial waters is consistent with the very low iron and manganese concentrations (Fig. F43) at Site 1256, as manganese and iron reduction occur in less strongly reducing conditions than does sulfate reduction. Iron concentrations are not given in Table T22 because concentrations are below the detection limit.
Dissolved silica in the interstitial waters is influenced by biogenic dissolution/reprecipitation reactions and, perhaps, basement alteration. Silica (Fig. F42) is released into the pore water during dissolution of diatom tests constructed of opal-A (Murray and Jones, 1992; Smith, 1998; Gieskes, 1974). This causes the increase in silica from a surface concentration of 483 µM to a maximum concentration of 1356 µM at 175 mbsf. Below 190 mbsf, dissolved silica decreases significantly to a minimum value of 111 µM at 226 mbsf and remains low (<300 µM) to basement. This decrease records the uptake into diagenetic opal-CT. In addition, the uptake of silica by chalcedony precipitation in the basement (see "Alteration" in "Basement Formed at Superfast Spreading Rate (Holes 1256C and 1256D)"), may also contribute to the low silica concentrations in the deepest pore waters.
The magnesium and calcium concentrations in interstitial waters exhibit opposite trends (Fig. F43) and show a strong inverse correlation (r = -0.97), as is observed at many DSDP and ODP sites. Concentration gradients of both elements are relatively constant (±0.24 mM/m) with depth, apart from an interval at 159 mbsf where concentrations change rapidly within 4 m. The dramatic chemical gradients across this interval (1.5 mM/m) indicate a layer of very low diffusivity, most probably a relatively impermeable chert bed. This interpretation is supported by the occurrence of a ~1-m-thick chert bed observed in the downhole measurements (see "Results" in "Downhole Logging" in "Downhole Measurements").
The strong inverse correlation between calcium and magnesium implies related behavior of these elements, typical of profiles influenced by diffusion between seawater and basement (McDuff and Gieskes, 1976; McDuff, 1981). The Mg/Ca ratio with depth (Fig. F44) does not, however, show a single gradient. There are at least three segments, without appreciable curvature, presumably reflecting changes in diffusive processes as a result of changes in the physical properties of the sediment or exchange of calcium and/or magnesium between water and sediment. Through the first segment (0-40.6 mbsf), Mg/Ca values linearly decrease from a near-seawater value of 5.2 to 3.2 at the lithologic Unit I/II boundary. Through the interval 40.6-159 mbsf, the ratios range from 3 to 2.2, showing a lower gradient than that observed in the upper 40 m of the sediment column. At 159 mbsf, the base of the impermeable chert layer, there is a distinct change toward lower ratios. Below 159 mbsf to basement, Mg/Ca ratios remain relatively constant, with a subtle decrease toward the basement. The Mg/Ca ratios of the near-basement samples are within the range of compositions measured in 30°-35°C basement fluids on the flanks of the Juan de Fuca Ridge (Mg/Ca = 1.0 to 1.5) (Elderfield et al., 1999).
Strontium (Fig. F43) generally increases in concentration toward basement and shows a great deal of variability. Strontium concentrations in interstitial waters can be strongly influenced by the formation of secondary carbonates by the dissolution and reprecipitation of biogenic carbonate and increase during alteration of volcanic material (Elderfield and Gieskes, 1982). Volcanic material is rare throughout the Site 1256 sediment column, and thus the latter process is probably of limited importance. Strontium is also only moderately correlated to calcium (r2 = 0.65). Sr/Ca ratios demonstrate a slight increase from seawater values with depth to a maximum (~12 mM/M) at 200 mbsf, but with some scatter throughout (Fig. F44). The values of the near-basement samples, however, are not consistent with what has been observed when interstitial waters are influenced by reaction with basalts. Basement fluids generally have greatly decreased Sr/Ca ratios because of large increases in calcium concentrations. On the Juan de Fuca Ridge flank, for example, 30°-35°C fluids have Sr/Ca ratios of ~3 mM/M (Elderfield et al., 1999), whereas at Site 1256, Sr/Ca ratios are two to three times higher. Therefore, other processes besides simple diffusion between seawater and basement are influencing the strontium concentrations of the interstitial fluids, most probably diagenetic reactions related to the recrystallization of biogenic carbonate, as is consistent with the low measured alkalinities.
Potassium concentrations decrease steadily downsection from ~11 mM at 2.2 mbsf to ~7 mM at 225 mbsf, with a slight increase above 8 mM near basement (Fig. F43). The total downhole change in potassium concentration is 36% and is similar to the trend of the magnesium profile. In fact, potassium and magnesium are inversely correlated (r = -0.92) and potassium and calcium are well correlated (r = 0.92). The K/Ca ratio shows a steady decrease with depth from seawater values to ratios typical of basement fluids (Elderfield et al., 1999) below 159 mbsf. K/Mg ratios are relatively uniform with depth (Fig. F44), suggesting similar behavior of potassium and magnesium throughout the entire sediment column. Below the chert layer at 159 mbsf, K/Ca ratios decrease because of a larger increase in calcium than potassium, displaying an average ratio of 0.33, similar to that measured in 30°-35°C basement fluids from the Juan de Fuca Ridge flank (Elderfield et al., 1999).
Lithium concentrations (Fig. F43) decrease from 22.1 µM at 2.2 mbsf to 11.4 µM at 159 mbsf, with increased concentrations between 75 and 95 mbsf. Below 159 mbsf, lithium concentrations decrease to 7.2 µM and then slightly increase to 8.47 µM at 216 mbsf. Below 216 mbsf, lithium concentrations greatly increase, with the bottommost sample having a concentration of 17.2 µM. From 75 to 95 mbsf, which is located in a lithologic unit of sandy silty nannofossil diatom ooze (see "Unit II" in "Description of Units" in "Lithostratigraphy"), the higher lithium concentration suggests an additional source, possibly the diatomaceous ooze. Below 216 mbsf, lithium concentrations increase toward seawater value, suggesting seawater contamination in Sections 206-1256B-26X-2 and 206-1256C-4R-1. This is reasonable because Section 206-1256B-26X-2 was split and sat for several hours before being sampled and Section 206-1256C-4R-1 consisted of small biscuits with limited means of trimming off all possible contamination. With the exception of these two anomalous intervals, the lithium profile is dominated by diffusion between seawater and basement.
The pattern of Li/Ca ratios is very similar to that of Mg/Ca ratios, suggesting the occurrence of diffusion between seawater and basement overprinted by diagenetic reactions (Fig. F44). Li/Mg ratios are similar to K/Mg ratios in that they are fairly uniform with depth; however, Li/Mg ratios display greater variability as well as possible contamination where there is a kickback toward higher seawater values below 216 mbsf.
The primary influence on the interstitial water chemistry at Site 1256 is diffusion between seawater and basement fluids. A chert bed located between 158 and 159 mbsf, observed in the downhole logs, demonstrates its low diffusivity by causing abrupt changes in cation concentrations, as seen in many of the depth profiles.
The low organic carbon content limits the extent of pore water sulfate reduction, as is also shown by decreases in alkalinity in the pore waters with depth. Alkalinity, conversely, reflects the dissolution and reprecipitation of biogenic calcite. Silica concentrations clearly illustrate the dissolution of biogenic opal-A during diagenesis and the precipitation of opal-CT. The concentrations of lithium and potassium mimic those of magnesium in the interstitial waters and are strongly influenced by diffusion between seawater and basement. This high input of biogenic silica, the low organic carbon content, and alkali element exchange with the sediments combine to demonstrate the extent of early and late diagenesis processes active at Site 1256.
The shipboard chemical analyses of sediment samples from Holes 1256B and 1256C included inorganic carbon, total carbon, nitrogen, sulfur, and major and trace element chemistry. Calcium carbonate (CaCO3) and organic carbon (Corg) concentrations were calculated from the inorganic carbon and total carbon analyses. Analytical methods and uncertainties are outlined in "Bulk Sediment Sampling and Chemical Analyses" in "Geochemistry" in the "Explanatory Notes" chapter). Data are reported in Table T23.
Calcium carbonate concentrations define four regimes with a distinct trend of increasing carbonate with depth, consistent with the lithologic units observed and changes in physical property data (see "Physical Properties") (Fig. F45). The uppermost 17.8 m of the sedimentary section shows an increase in carbonate from a few weight percent to ~43 wt%, with some variation, reflecting the silty clay and clay-rich nannofossil ooze in lithologic Subunit IA. From 17.8 to 35 mbsf, calcium carbonate values are low, averaging ~8 wt% within the sandy silty clay that composes lithologic Subunit IB (see "Description of Units" in "Lithostratigraphy"). From ~35 to 123 mbsf, which includes the base of lithologic Subunit IB, calcium carbonate concentrations increase, averaging ~30 wt%, although anomalously low carbonate concentrations were measured at 89.5 mbsf. The boundary between lithologic Units I and II is not clearly defined by calcium carbonate concentrations. At ~123 mbsf, calcium carbonate concentrations increase to ~85 wt% and remain uniformly high down to the lowermost sample at 245 mbsf, with the exception of a single low carbonate concentration at 173 mbsf. The average concentration in this interval is 80 wt%, which is typical of carbonates in the equatorial Pacific. The concentrations of calcium carbonate measured are also consistent with the visible and near-infrared spectroscopy (VNIS) data (see "Visible and Near-Infrared Spectroscopy" in "Physical Properties").
Total carbon concentrations indicate that most of the carbon is inorganic (Table T23). Total organic carbon values (Fig. F45), calculated as the difference between total carbon and inorganic carbon, are low throughout the sediment column, ranging between 0.03 and 2.4 wt% (average < 1 wt%). Total organic carbon concentrations average 0.6 wt% in the surface sediments, although most of the sediments downhole contain <0.25 wt%, a trend typical of open ocean sediments in this region (Lyle, 1992). A single measurement from 173 mbsf yields a Corg of 8.9 wt% and coincides with a very low calcium carbonate content. Elevated Corg concentrations (2-2.4 wt%) were also measured in a few samples from Cores 206-1256-25X and 26X at 216-227 mbsf, but a strong trend is not exhibited. Total nitrogen and sulfate values at Site 1256 are very low and are <0.14 and <0.45 wt%, respectively. In most of the sediment samples, nitrogen and sulfur contents are too low to be measured accurately. As such, organic C/N ratios were not used to determine the origin of the organic matter.
Bulk sediment chemistry reflects changes in the relative contribution of biogenic silica, biogenic carbonate, and terrigenous material to the sediments. There is a chemical transition at ~115 mbsf, at the base of a ~4-m-thick strongly laminated diatom mat. The overall elemental compositions indicate clay-rich and/or silica-rich sediments in the upper 115 mbsf and dominantly carbonate-rich sediments below. There is also a distinct peak at ~187 mbsf, with higher concentrations of silica, titanium, aluminum, iron, manganese, magnesium, sodium, and potassium indicating an interval with a high terrigenous content.
Silica concentrations (expressed as weight percent oxide) define the two sediment types (Fig. F46). Above 115 mbsf silica concentrations average 58 wt%, whereas below this level concentrations steadily decrease from roughly 28 to 8 wt%. The silica profile is consistent with the higher abundance of detrital components in the upper 100 m and their general absence below. Siliceous microfossils are of significant but variable importance throughout the sedimentary section except directly above basement; this is reflected in low silica concentrations from 200 to 250 mbsf. Chert nodules and layers were recovered from 111 mbsf to basement, and the presence of these diagenetic features complicates the interpretation of downhole silica concentrations.
Calcium concentrations (Fig. F46) increase from an average of 3 wt% oxide in the upper 20 m to ~17 wt% above the 115-mbsf transition. Below 115 mbsf, calcium concentrations increase significantly and remain relatively uniform (45 ± 5 wt%) with the exception of anomalously low carbonate contents at 173 mbsf (11 wt%) and 187 mbsf (35 wt%). The calcium profile mirrors the silica profile, and these elements are inversely correlated (r = -0.95).
Aluminum, titanium, iron, magnesium, sodium, and potassium concentrations (expressed as weight percent oxides) (Fig. F46) all decrease from 2.2 to 115 mbsf, at which point there is a further abrupt decrease in concentrations. Below 115 mbsf, these concentrations remain low with limited variability to basement, with the exception of an interval at 187 mbsf where slightly higher concentrations are recorded. Aluminum, titanium, iron, magnesium, sodium, and potassium have very similar downhole variations.
Aluminum concentrations can be used to estimate the abundance of terrigenous material in the sediment column (Murray et al., 1999), where,
This equation compares measured aluminum concentration to that of average post-Archean average shale (PAAS; Al = 10 wt%) (Taylor and McLennan, 1985) to assess the relative proportion of terrigenous material in the sediments. PAAS may not be exactly appropriate for comparison with the terrigenous material in the Site 1256 sediments but does allow comparison with other sedimentary sequences.
Terrigenous material is most abundant in the uppermost 40 m and decreases steadily to 115 mbsf, below which terrigenous components compose <5% of the sediment (Fig. F45). Terrigenous material accounts for 20% of the sediment in the anomalous sample at 187 mbsf.
The percentage of biogenic silica can be estimated from the relative proportions of silica and aluminum in the sediments. The crustal ratio of SiO2/Al2O3 is roughly 3.0 (Taylor and McLennan, 1985), and previous measurements of this ratio in Cenozoic sediments from the equatorial Pacific Ocean (Leinen, 1979) indicate that this value should be reasonable for Site 1256. Therefore, an estimate of the biogenic silica content of the sediments (Fig. F45) can be obtained using the following equation:
The uncertainty using this equation is estimated to be approximately ±5 wt% biogenic silica (Leinen, 1977).
Using this formula to calculate the proportion of biogenic silica indicates that the Site 1256 sediments contain ~20-50 wt% biogenic silica in the upper ~111 m and that biogenic silica decreases from ~20 to 5 wt% below 115 mbsf. Between 111 and 115 mbsf, in the interval of the diatom mat, biogenic silica composes 74 wt% of the sediment (see "Lithostratigraphy"). Distinct chert horizons were observed in the downhole logging data at 111 and 158 mbsf (see "Results" in "Downhole Logging" in "Downhole Measurements"), and chert nodules are common throughout the lower part of the sedimentary section. The chert nodules were not sampled for chemical analysis.
Lastly, Fe/Al ratios can be used to indicate the metalliferous component of marine sediments. The Fe/Al value for average shale (PAAS) is ~0.43 (Taylor and McLennan, 1985), and Site 1256 sediments range from 0.78 to 11.9. Fe/Al ratios remain relatively constant, with an average value of 0.79 in the upper 66 m of the sediment column. From 66 to 159 mbsf, ratios gradually increase from 1.6 to 1.8 with two anomalous values of >2 at 112 and 123 mbsf. Below 159 mbsf, Fe/Al values increase from 2.3 to 6 with depth and at ~243 mbsf display an anomalous ratio of 11.9. In general, Fe/Al ratios indicate a significant metalliferous component in the sediments that increases with depth.
In summary, the sediment deposited at Site 1256 is predominantly composed of calcium carbonate, terrigenous material, and biogenic silica with an influence of a metalliferous component downhole (Fig. F45). Calcium carbonate is dominant below 115 mbsf, although at 187 mbsf the terrigenous component increases to 20%. At ~115 mbsf, biogenic silica is the most abundant phase. Above 115 mbsf, the proportion of terrigenous material increases toward the surface and is the most important component in lithologic Unit I (above 40.6 mbsf). These observations reflect changes in the depositional environment with time.
To illustrate changes in depositional environments through time and the processes that produced the sedimentary record at Site 1256, mass accumulation rates (MARs) (expressed as grams per square centimeter per 1000 yr) have been calculated for the major components of the sediments. We have calculated MARs for calcium carbonate, terrigenous material, biogenic silica, and metalliferous component from the weight percent of each sediment type, the linear sedimentation rates (LSRs) estimated from paleomagnetic and microfossil constraints (in centimeters per 1000 yr) (see "Sedimentation Rates"), and the measured bulk density (BD) of the sediments (in grams per cubic centimeter) (see "Density" in "Physical Properties"). For example, the CaCO3 mass accumulation rates were calculated as follows:
Mass accumulation rates of calcium carbonate show a clear transition at ~10.8 Ma (~115 mbsf), where there is a significant change in the nature of carbonate deposition. During the middle Miocene, carbonate MAR is high with the exception a major decrease at 12.5 Ma (~173 mbsf). There is also a sharp reduction at 10.8 Ma (~115 mbsf) within the diatom mat. Carbonate MAR is extremely low from ~10.2 to 8.4 Ma, with carbonate deposition virtually ceasing at the end of the middle Miocene (~9.6 Ma). Carbonate sedimentation remained low (<1.5 g/cm2/k.y.) from the late Miocene to present, with the accumulation of carbonate particularly low (<0.2 g/cm2/k.y.) from 5 to 2 Ma.
The nature of carbonate deposition and preservation at Site 1256 has been influenced by a number of processes that have all led to a reduction in carbonate accumulation. Site 1256 formed on the eastern flank of the EPR at ~15 Ma at an equatorial latitude (see Fig. F11 in the "Leg 206 Summary" chapter) within the zone of high equatorial productivity. Superfast spreading away from the axis has transported Site 1256 northeastward out of the zone of high equatorial productivity. The movement of the crust away from the ridge axis will also result in a deepening of the seafloor with age, with the crust potentially subsiding below the CCD. Consequently, the deposition and preservation of calcium carbonate is reduced. These two processes should result in a gradual decrease in carbonate mass accumulation; however, more severe changes are observed. The significant changes in the preservation and production of calcium carbonate may also have been from rapid shoaling of the CCD as a result of changing deepwater chemistry or large changes in oceanic productivity. It is difficult to decipher the relative importance of these mechanisms at this preliminary stage of investigation.
This major decrease in carbonate MAR was also observed in sedimentary records drilled during Leg 138, and this event is referred to as the carbonate crash (Lyle et al., 1995; Farrell et al., 1995). Their measurements of carbonate concentrations and calculations of MARs from various sites in the central and eastern equatorial Pacific show a very large decline in calcium carbonate and a large decrease in carbonate MAR near the middle/late Miocene boundary (9.5 Ma). Farrell et al. (1995) demonstrated that this event began in the Guatemala Basin between 12.5 and 12 Ma, followed by fluctuations in opal productivity and enhanced dissolution of carbonate that ended in the most intense reduction in carbonate accumulation at 9.5 Ma. This event has also been observed in the Caribbean Sea at Site 998 based on the poor preservation of calcareous fossils. The timing of the Caribbean event is 1-2 m.y. earlier (Sigurdsson, Leckie, Acton, et al., 1997). The possible cause of the carbonate crash has been suggested to be the constriction of the Panama Gateway and/or the onset of North Atlantic Deep Water formation (Lyle et al., 1995; Farrell et al., 1995).
Mass accumulation rates of terrigenous material (Fig. F47) at Site 1256 are low during the middle Miocene with two substantial peaks at 12.9 Ma (~189 mbsf) and 10.5 Ma (~104 mbsf), slightly above the diatom mat. After 8 Ma, during the late Miocene, the accumulation of terrigenous material increased with increased MARs between 7 and 5 Ma and from 1.8 Ma to present. These elevated MARs are consistent with the abundance of eolian material observed in Leg 138 sedimentary records. Eolian grain-size variability and flux records from Leg 138 sites suggest that the strongest southeast trade winds occurred between 8 and 5 Ma, reflecting the position of the Intertropical Convergence Zone and/or latitudinal position of Site 1256 through time. Increased terrigenous accumulation from 1.8 Ma to present may be a response to the increased intensity of the northeast trade winds at ~1 Ma, resulting from the onset of Northern Hemisphere glaciation (Hovan, 1995).
Mass accumulation rates of biogenic silica (Fig. F47) show a slight increase with some variation from the bottom of the hole to the middle Miocene (10.8 Ma). There is a peak in accumulation at 10.8 Ma, where the diatom mat is located. In sediments younger than 10.8 Ma, MARs are low (<0.5 g/cm2/k.y.) except for elevated rates between 5 and 8 Ma. Changes in biogenic silica MARs can reflect changes in productivity as well as preservation, where dissolution occurs because of pronounced undersaturation of seawater with respect to silica (e.g., Hurd, 1973, and many others). The increased biogenic silica MAR between 5 and 8 Ma may record a globally observed biogenic bloom (Farrell et al., 1995). Site 1256 is similar to other eastern equatorial sites where a biogenic bloom occurred between 4.5 and 6.7 Ma, evidenced by extremely high MARs (Pisias et al., 1995). This biogenic bloom has been linked to higher productivity, but the mechanisms are still uncertain.
Mass accumulation rates of the metalliferous component show a significant accumulation (~0.8 g/cm2/k.y.) from ~14.5 to ~14.2 Ma. After ~14 Ma, MARs decrease significantly to ~0.15 g/cm2/k.y. and continue to decrease to <0.05 g/cm2/k.y. at ~10.2 Ma. These rates remain low to the present day. This decreasing trend reflects the plate movement off the ridge of the EPR.
Ba/Ti ratios are often used as a chemical proxy of the biogenic productivity in the overlying waters, since sedimentary calcium carbonate, opal, and organic carbon are diagenetically labile and it is difficult to directly link their final concentrations in the sediment to variations in productivity of the overlying water. Titanium is largely unaffected by diagenetic remobilization, and barite (BaSO4) is extremely insoluble, so primary Ba/Ti ratios should be preserved during deposition and diagenesis (e.g., Murray et al., 1999). Ba/Ti ratios are a useful proxy of biogenic flux because barite precipitates in close association with settling biogenic matter in regions of high biogenic productivity, although many details of the biogeochemical linkages between barium and export production remain unclear (Dymond et al., 1992; Francois et al, 1995). The variations in barite abundance and, by inference, biogenic production can be measured by the bulk chemical analysis of barium through the sediment column. The Ba/Ti ratio is particularly useful for Site 1256 because sulfate reduction, as seen in the interstitial water analyses, is minimal (SO42- >18 mM) and barite is unlikely to have been significantly altered during diagenesis.
Ba/Ti ratios increase from ~2 to ~10 in the upper 112 m of Site 1256, and values below 112 mbsf are relatively constant with an overall average of 23 ± 6 (Fig. F48). These Ba/Ti ratios are much higher than that of average shale (Ba/Ti = 0.11) (Taylor and McLennan, 1985), and the values below 112 mbsf are within the range of 20-30, typical of the eastern equatorial Pacific Ocean (Schroeder et al., 1997). There is a significant decrease in biogenic production following an extended episode of high productivity between 14.6 and 10.8 Ma. However, at 12.9 Ma, Ba/Ti ratios drop below 10, reflecting a very brief period of low productivity. More significantly, productivity decreased significantly at 10.8 Ma and has remained low to modern day. This change in productivity could be a response to the northward movement of the Cocos plate away from the equator, where productivity and, hence, Ba/Ti ratios are commonly higher. Ba/Ti ratios do not directly correspond to MARs of biogenic silica and do not indicate a significant increase in productivity between 5 and 8 Ma.
In summary, sedimentary data from Site 1256 are consistent with other eastern equatorial Pacific sites. Bulk sediment data show a decline in calcium carbonate deposition at ~10.8 Ma with a carbonate crash most extreme at 9.6 Ma. Mass accumulation rates of biogenic silica before the carbonate crash show an increasing trend, with the exception of a drop at 12.8 Ma. After the carbonate crash, silica MARs are low with slightly elevated rates between 8 and 5 Ma, possibly associated with a global biogenic bloom. The proportion of terrigenous material increases gradually with time after the carbonate crash and displays significantly higher accumulation rates between 8 and 5 Ma and from 1.8 Ma to present.
We characterize the physical properties of the sediments cored in Holes 1256A, 1256B, and 1256C through a series of measurements on whole-core and split-core sections and discrete samples as described in "Physical Properties" in the "Explanatory Notes" chapter. In addition to the standard suite of physical properties measured during ODP cruises, which includes magnetic susceptibility, density, porosity, P-wave velocity, natural gamma ray (NGR) activity, and thermal conductivity, we also use light reflectance properties in the visible and near-infrared wavelengths to estimate the bulk composition of the sediment in terms of percent clay, silica (opal), and calcite. Taken together, these observations indicate that changes in physical properties are mostly gradual downhole, with the exception of those properties that are sensitive to major lithologic and compositional changes, which occur at the lithologic Unit I/II boundary, at the diatom mat (111-115 mbsf), and ~40-45 m above basement (~205-210 mbsf), below which chert and glauconite increase. One slightly atypical feature of the physical properties is how little the P-wave velocity varies downhole. Instead of increasing gradually downhole, as might be expected with increasing compaction, the velocity is nearly uniform at ~1540 m/s from 0 to ~205 mbsf. These and other features of the physical properties at Site 1256 are described below.
Whole-core sections were measured on the MST. The measurements consisted of GRA bulk density, NGR activity, and magnetic susceptibility on whole-core APC and XCB sections, plus P-wave logger (PWL) measurements on APC sections (Figs. F49, F50). Measurement spacing was 2.5 cm for GRA and magnetic susceptibility and 5 cm for the PWL and NGR activity, with an NGR measurement time set at 15 s (see "Multisensor Track Measurements," in "Physical Properties" in the "Explanatory Notes" chapter).
In general, the quality of the measurements was high, with the usual exception of intervals where the core did not totally fill the core liner or where drilling disturbance had occurred (e.g., soupy intervals). Data that were clearly erroneous, such as spikes or other extreme outliers, were excluded prior to interpretation. All P-wave velocity measurements from 0 to 63.15 mbsf made by the PWL were in error because the housing on the distance transducer was loose, resulting in erroneously long distance measurements that in turn resulted in velocities that were faster than the true velocities. Even after the distance transducer was repaired, the data from the PWL were anomalous. They show a relatively large increase in velocity downhole from ~1511 m/s at 64 mbsf to >1600 m/s at 122 mbsf, which disagrees with the discrete measurements made on the split-core sections. Moreover, this increase mimics that seen in the upper 63 mbsf where the distance transducer became loose. Thus, we do not use the PWL data in the interpretations below.
Thermal conductivity measurements were made at a frequency of one to three per core. Moisture and density properties (porosity, bulk density, and grain density) were measured on one sample per section for most cores from Holes 1256A and 1256B (Figs. F49, F50). P-wave velocities (Hamilton Frame PWS3 contact probe) (Fig. F49) were measured once per section on split cores, with a correction made for the core liner.
The GRA bulk density from the whole-core sections and the bulk density from the discrete samples are in good agreement (Fig. F49). The grain density follows approximately the same general trend but at a higher density, as is expected. The open symbols in the interval from 26 to 42 mbsf on Figure F49 are data points that fall outside the normal range of values. The anomalous values in bulk density, grain density, and porosity are associated with the same set of discrete samples that were run in two different batches in the pycnometer. It is not clear if the anomalous values result from instrument error or some other error in handling, but clearly they are in error, since in several of the samples the grain density is even lower than the bulk density. Not including the above-mentioned samples, the bulk density averages ~1.15 g/cm3 from 0 to 96 mbsf. From 96 to 111 mbsf, which is just above the diatom mat, densities are slightly higher (~1.25 g/cm3) and more variable. The low-density trough (~1.12 g/cm3) at ~111 mbsf marks the diatom mat, below which the density steadily increases to ~1.45 g/cm3 at 125 mbsf and then very gradually increases downhole to ~1.7 g/cm3 at the basement contact. Grain densities are more variable and slightly lower from the mudline to the base of the diatom mat (0-115 mbsf) than for the sediment below. The average grain densities above and below the diatom mat are 2.42 and 2.67 g/cm3, respectively.
Porosity data display a trend that is roughly the inverse of the density profile (Fig. F49). Ignoring the outliers (open symbols in Fig. F49), the porosity gradually decreases downhole from a high of ~90% near the mudline to a low of 55% near the basement, with a step decrease of ~20% across and slightly below the diatom mat from 111 to 125 mbsf.
P-wave velocity from discrete samples does not show an increase in velocity with depth (Fig. F49), which is at least a bit atypical for a sedimentary section more than 200 m thick. The velocity is nearly constant at an average value of ~1540 m/s from 0 to ~205 mbsf. The only notable anomaly occurs at the diatom mat (111-115 mbsf), where there is a sharp downhole increase in velocity at the top of the diatom mat with velocity decreasing back to ~1540 m/s down to ~205 mbsf. Below 205 mbsf, the velocity increases but is variable, with values ranging between 1520 and 1760 m/s.
Except for sediments below 205 mbsf, the velocity does not increase with the decrease in porosity that occurs downhole, counter to what would be predicted by Gassmann's (1951) theoretical model for the controls on velocity in porous rocks. Porosity is not explicitly included in this model, but porosity implicitly affects velocity through its influence on frame bulk modulus (Hamilton, 1971), shear modulus (Stoll, 1989), and bulk density. Consequently, siliciclastic rocks and sediments from all parts of the world exhibit a strong dependence of P-wave velocity on porosity (e.g., Wyllie et al., 1958; Erickson and Jarrard, 1998). Instead, for most of the sediment, velocities actually increase slightly at higher porosities (Fig. F51).
NGR activity is relatively high at the mudline and then sharply decreases from the high of 36 counts per second (cps) to 15 cps at ~9 mbsf (Fig. F50). It then decreases downhole very gradually, reaching a near-background average of 13 cps just above basement. The deviations from this otherwise smooth trend are (1) a minor trough from 38 to 45 mbsf, possibly related to the lithologic Unit I/II boundary; (2) a minor increase from 73 to 79 mbsf, possibly a result of a layer of sandy silt; and (3) a more variable zone from just above the diatom mat to the base of the mat (108-115 mbsf).
Magnetic susceptibility is characterized by three main subdivisions. From 0 to 40.6 mbsf (lithologic Unit I), the susceptibility is high (average = 28.7 raw susceptibility meter units), owing to the high clay content of the sediments. From 40.6 mbsf to the top of the diatom mat there is a transition zone as the sediments become more calcareous and the susceptibility decreases to near zero. Between the top of the diatom mat and ~205 mbsf, the susceptibility is very low, often giving 0 or negative values in raw susceptibility meter units. Below ~205 mbsf the magnetic susceptibility increases to an average of ~7.5 raw meter units, an increase that correlates with the increase in glauconite and chert observed in this interval. These trends, their relationship to other magnetic measurements, and comparisons between whole-core and point susceptibility measurements are discussed further in "Paleomagnetism".
Thermal conductivity values are nearly uniform at ~0.7 W/(m·K) from 0 to 111 mbsf (Fig. F50). A step increase occurs at the top of the diatom mat at 111 mbsf to a value of ~1.0 W/(m·K). This value is maintained all the way to ~210 mbsf with only slight variation. At 210 mbsf, thermal conductivity increases slightly downhole, reaching a high of 1.2 W/(m·K) just above basement. For the effect of thermal conductivity on heat flow, refer to "Temperature and Heat Flow" in "Downhole Measurements."
VNIS measurements were undertaken on one sample per section throughout the sedimentary section at Site 1256. Measurement methods are described in "Visible and Near-Infrared Spectroscopy" in "Physical Properties" in the "Explanatory Notes" chapter.
VNIS measures the spectrum of light reflected from a rock surface. Light is absorbed by minerals at and near the rock surface by both electron and vibrational processes (Clark et al., 1990; Clark, 1995). Types of identifiable minerals depend on the frequency band of the instrument. We use the FieldSpec Pro FR portable spectroradiometer because of its wide bandwidth (350-2500 nm). At near-infrared wavelengths (950-2500 nm), normal modes of characteristic vibrations of OH bonds occur; thus, diagnostic absorption bands for water (H-OH), Mg-OH, Al-OH, and Fe-OH are evident.
VNIS has a sufficiently broad detection band (350-2500 nm) for identification of the most abundant minerals in equatorial Pacific sediments: opal (SiO2), calcite (CaCO3), and clay (calibrated for smectite during Leg 199). During Leg 199, VNIS was used to provide preliminary percentages of these minerals for ~1000 samples (Shipboard Scientific Party, 2002; Vanden Berg and Jarrard, 2002a); postcruise analyses are refining those mineral percentages (Vanden Berg and Jarrard, 2002b). The spectral analysis algorithm of Vanden Berg and Jarrard (2002b) has been used to convert spectra from Site 1256 sediments to preliminary mineral percentages (Fig. F52; Table T24). These estimates are expected to be revised postcruise, based on calibration with additional standards. Comparison of VNIS-based calcite estimates with coulometer calcite measurements (Fig. F52) shows that VNIS successfully captures the overall pattern of variability within Site 1256, although some systematic offsets are evident (Fig. F53). For example, predicted carbonate values are consistently high for low (<20 wt%) and high (>85 wt%) measured carbonate and are consistently low when the measured carbonate falls between 40 and 80 wt%. Fine-tuning the VNIS measurement to the Site 1256 sediments will undoubtedly lead to more accurate predicted compositions.
During coring of Hole 1256B, several temperature measurements were taken with the APCT tool for the purpose of measuring heat flow. Temperature readings were successfully taken in bottom water prior to Core 206-1256B-1H and in sediment prior to retrieving Cores 6H, 9H, 13H, and 17H. A reading for Core 206-1256B-4H was unsuccessful because of apparent battery failure. Temperatures increased from 1.4° to 25.0°C, with the thermal gradient decreasing downhole from 0.165°C/m to 0.117°C/m (Table T25). Heat flow is the product of the thermal gradient and the thermal conductivity. With the conductivity increasing downhole (see "Thermal Conductivity" in "Physical Properties"), the heat flow decreases more slowly than the thermal gradient, from 120 mW/m2 at 0-54 mbsf to 109 mW/m2 at 120-158 mbsf (average = 113 mW/m2) (Fig. F54). Assuming that the deepest measured heat flow of 109 mW/m2 remains uniform to the base of the sediment section, the average thermal conductivity of 1.03 W/(m·K) below Core 206-1256B-17H implies a temperature of ~35°C at the sediment/basement interface.
The value of 113 mW/m2 is close to that predicted by models of conductive cooling of oceanic lithosphere (Parsons and Sclater, 1977; Stein and Stein, 1992), implying that hydrothermal circulation is no longer a major mechanism of heat transport at Site 1256. This area has been previously noted for its low heat flow (Langseth and Silver, 1996); other measurements in the area are typically 10-70 mW/m2 (Louden and Wright, 1988). The thick ponded lava flow in the upper basement may be a factor in restricting circulation here relative to nearby areas with comparable sediment cover, both by providing an impermeable cap to the basement and by limiting topographic relief, which might help drive circulation within the sediment. The downhole decrease in heat flow appears to be slightly above uncertainty. If indeed real, a slight reduction in bottom water temperature over the last few thousand years seems more plausible than upward advection of pore water as an explanation for the trend.
Logging operations were conducted in Hole 1256C after it had been drilled ahead to a depth of 220 mbsf with a 9.75-in drill bit and cored from this depth to 331 mbsf. At the conclusion of coring, the hole was prepared for logging with a wiper trip to 83 mbsf. After the bit was released, the hole was displaced with 92 bbl of sepiolite mud and the EOP was set at 125.8 mbsf; during logging, the pipe was raised 17 m. Three tool string runs were planned: the triple combo, FMS-sonic, and BGR magnetometer. Operational difficulties resulted in data being collected only by the triple combo tool string.
The triple combo tool string was deployed first. It included, from top to bottom, the Hostile Environment Gamma Ray Sonde (HNGS), the Accelerator Porosity Sonde (APS), the Hostile Environment Litho-Density Tool (HLDT), the Dual Laterolog (DLL) tool, and the Lamont-Doherty Earth Observatory (LDEO) Temperature/Acceleration/Pressure (TAP) tool. This tool string was initially unable to pass a bridge at ~203 mbsf. After several unsuccessful attempts to pass the tool string through this bridge, only the upper part of the borehole was logged. Two logging runs were performed from this depth to the EOP. After these passes, the tool string was recovered and set back in the derrick. The drill string was then lowered to within 20 m of the bottom of the hole to clear obstructions, and another attempt was made to deploy the triple combo tool string with the EOP this time placed at 231 mbsf; during the run the pipe was raised 16 m. During this third run, basement and overlying sediment were logged. The FMS-sonic tool string was then deployed but was unable to pass deeper than 257 mbsf (26 m below the EOP and only 5 m into basement). After several unsuccessful attempts to pass the tool string through this obstruction, it was decided to terminate the logging program.
The wireline depth to seafloor is usually determined from a step increase in gamma ray values at the sediment/water interface. Unfortunately, because of the extremely low gamma ray values of the sediments, this depth could not be determined in Hole 1256C. Instead, we referenced the wireline depth measurement to the pipe depth measurement (mbsf scale) at the sediment/basement interface, where a clear decrease occurs in gamma ray activity along with a step increase in electrical resistivity. The wireline reference was then applied to all three runs. Small differences exist in wireline measurements and pipe measurement. For example, the obstruction at 203 mbsf was possibly as deep as 207 mbsf based on the wireline measurement. The pipe measurement in this case was only intended as a relative crude estimate for where the driller felt a constriction and may not be identical to the depth of the bottom of the triple combo tool string.
Degraded borehole width affects most measurements, particularly those that require eccentralization and good contact between the tool and the borehole wall (APS and HLDT). In the first logged interval from ~203 mbsf (corresponding to ~207 m wireline depth) to the EOP, the caliper remained fully open (16.5 in, which is the maximum range of the tool), except between 152 and 162 mbsf, where a small constricted interval was recorded by the caliper (13.8 in). Poor borehole conditions along the entire logged interval resulted in poor-quality HLDT measurements, as evidenced by logs that have erroneous spikes in the density and photoelectric factor (PEF). During the third logging run, 22 m of sediment above the basaltic basement was logged. In this interval caliper data range from 12 to 17 in, with only a narrow washed-out spot at 233 mbsf. Consequently, in this interval the density probe provided reliable measurements without any erroneous spikes. Resistivity logs are relatively insensitive to borehole size, and the NGR logs were corrected for borehole size during acquisition. Despite the washed-out conditions, gamma ray, porosity, and resistivity data from the upper part of the borehole are of good quality. No TAP tool data could be recovered, probably due to a power failure during the long (24 hr) triple combo run.
Results from both tool string passes are presented in Figure F55. For the lower pass, only the sediment and the sediment/basement interface are discussed here; the measurements recorded in the basement are presented in "Downhole Measurements" in "Basement Formed at Superfast Spreading Rate (Holes 1256C and 1256D)." Operational difficulties (hole obstructions) resulted in discontinuous data. Furthermore, with the triple combo tool string used in Hole 1256C being 41.3 m long, the relative positions of the probes on the tool string result in data being recorded over different intervals (see the "Downhole Measurements" in the "Explanatory Notes" chapter).
Over the entire sedimentary section, natural radioactivity appears to be low, ranging from 4 to 10 gAPI (American Petroleum Institute units) with a slight decrease below 165 mbsf. In the lower logged interval, a slight increase in natural radioactivity is recorded from 220 mbsf to the sediment/basement interface at 251 mbsf. In this lower interval, the potassium content significantly increases (from 0.5 wt% in the upper interval to 2 wt% in the lower interval); this increase can be related to glauconite in the recovered core material (see "Lithostratigraphy"). Five distinctive peaks (up to 20 gAPI) in natural radioactivity are recorded in the upper section at 111, 122, 129, 136, and 158 mbsf. Two of these peaks (111 and 158 mbsf) are associated with a strong decrease in porosity (from 75% down to 25%). The peak at 158 mbsf is also associated with a strong increase in electrical resistivity. These two intervals may correspond to chert layers. The peak in natural radioactivity at 122 mbsf is associated with a strong increase in thorium content and could correspond to an ash layer.
The porosity measurements from the logs are higher than the measurements made on discrete core samples (see "Porosity" in "Physical Properties") because the neutron porosity log responds to hydrogen in the formation and also measures bound water and fracture porosity (Schlumberger, 1989; Rider, 1996). In all, porosity measurements show a gentle decrease downsection, illustrating the compaction sustained by the sediments.
Overall, the logged intervals appear to be relatively homogeneous. The sedimentary sequence below 110 mbsf consists predominantly of nannofossil ooze with minor intervals of biogenic silica-rich sediment. The logs show no obvious trends with depth other than a step change at the contact between the sediment and the basaltic basement, which is marked by a sharp increase in resistivity and density and a corresponding decrease in porosity. The electrical resistivity shows a strong increase from 1.2 
m in the sediments to 28 m in the basaltic basement. Porosity similarly recorded the transition with a sharp decrease from 60% in the sediments to 10% in the basement.
