Questioning carbonate diagenetic paradigms: evidence from the Neogene of the Bahamas

Transcription

1 Marine Geology 185 (2002) 27^53 Questioning carbonate diagenetic paradigms: evidence from the Neogene of the Bahamas L.A. Melim a;, H. Westphal b, P.K. Swart c, G.P. Eberli c, A. Munnecke d d a Western Illinois University, Macomb, IL 61455, USA b Institut fu«r Geologie undpala«ontologie, Universita«t Hannover, Callinstr. 30, Hannover, Germany c RSMAS, University of Miami, 4600 Rickenbacker Cswy., Miami, FL 33149, USA Institut fu«r Geologie undpala«ontologie, Universita«t Tu«bingen, Herrenberger Str. 51, Tu«bingen, Germany Received 8 May 2000; accepted 20 September 2001 Abstract Carbonate diagenetic models have been heavily influenced by numerous studies of exposed Quaternary limestones. As a result, meteoric diagenesis is often assumed to be the principle means of altering aragonite-rich sediments into calcitic limestones. However, these models are limited by the scarcity of examples of aragonite-rich sediments buried in seawater that have never been influenced by meteoric fluids. The Bahamas transect cores recovered originally aragonite-rich sediments deposited in deep water beyond the easy reach of meteoric waters and provide an opportunity to test current diagenetic paradigms. The Bahamas transect consists of seven cores drilled in the prograding western margin of Great Bahama Bank. The two proximal cores (Clino and Unda) were drilled on the platform top and recovered shallow-water platform to reef facies overlying deeper margin and proximal slope facies. The five distal cores were drilled by ODP Leg 166 in up to 660 m of water and recovered carbonate slope facies. All studied sections are Neogene to Pleistocene in age. Diagenetic environments were identified based on petrographic and scanning electron microscopy (SEM) observations, XRD mineralogy, carbon and oxygen stable isotopic data, and trace elements. The upper 100^150 m of the two proximal cores were altered in meteoric to mixing-zone diagenetic environments but all other intervals were altered exclusively in marine pore fluids during seafloor, marineburial, and deep-burial diagenesis. Several of the findings of this study question current carbonate diagenetic paradigms. These include: (1) large-scale sea level lowstands may not have chemically active meteoric lenses as we found no meteoric alteration at the 3120 m elevation of the latest Pleistocene lowstand. Rather, phreatic meteoric diagenesis appears restricted to within W10 m of the land surface. (2) Mixing-zone diagenesis includes aragonite dissolution and minor LMC cementation but does not show the cavernous porosity or dolomitization predicted by mixing-zone diagenetic models. Current models are based on coastal mixing zones, which do not appear to be applicable to these more inland, and perhaps more typical, locations. (3) Marine-burial diagenesis produces a mature limestone with fabrics formerly considered diagnostic for meteoric diagenesis such as moldic porosity, aragonite neomorphism, blocky calcite spar and calcite microspar. However, oxygen stable isotopic data (average N 18 O=+1x) indicate alteration in marine pore fluids only. The character of marine-burial diagenesis is partially controlled by the nature of the sediment being altered. We have identified two end-member styles, an open-system style characterized by dissolution of aragonite without significant cementation and a more closed-system style with aragonite dissolution * Corresponding author. address: la-melim@wiu.edu (L.A. Melim) / 02 / $ ^ see front matter ß 2002 Elsevier Science B.V. All rights reserved. PII: S (01)

2 28 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 accompanied by calcite cementation. The sediments examined were deposited well above the aragonite compensation depth, so seawater entering the sediment is saturated with respect to aragonite. The under-saturation needed to drive diagenesis is likely the result of bacterial oxidation of organic matter using sulfate. (4) Microspar forms in these sediments as a cement based on petrographic and SEM examination of partly to completely altered samples. This contradicts the common assumption that microspar forms by aggrading neomorphism of micrite. (5) Strontium content of sediments altered in marine pore fluids can show an extreme range of values, formerly thought to indicate different environments. The opportunity to finally examine the diagenesis of aragonite-rich sediments buried in seawater challenges current diagenetic paradigms and emphasizes the importance of integrated studies. ß 2002 Elsevier Science B.V. All rights reserved. Keywords: Bahamas; carbonate diagenesis; marine-burial diagenesis; meteoric; microspar; calcite cement; mixing zone; strontium; aragonite 1. Introduction Previous work on the diagenesis of metastable carbonates (aragonite and high-mg calcite) has focused mainly on surface sediments from shallow marine tropical environments (for a review see James and Choquette, 1983, 1990b). Because Neogene to Quaternary sea-level uctuations have led to meteoric in uence on most young shallow-water sediments, carbonate diagenesis has long been thought to be dominated by meteoric alterations. More recent research, however, has demonstrated that the conditions under which those sediments have been altered are neither representative for the all of Earth s history, nor for deeper-water tropical settings (for an overview see Bathurst, 1993). Periplatform carbonates adjacent to modern tropical carbonate platforms are characterized by high fractions of bank-derived aragonite and high-mg calcite (James and Choquette, 1983). Prior to 1985, when ODP Leg 101 took place, most information on carbonate diagenesis of platform slopes was drawn from piston cores (e.g. Schlager and James, 1978; Mullins, 1983, 1986; Mullins et al., 1985). Studies of Saller (1984) and Schlager et al. (1988) are some of the few earlier investigations based on deep cores. In 1985, ODP Leg 101 o ered a rst opportunity to study diagenetic alterations of periplatform sediments in deeper cores from the lower slope to toe-of-slope of the Bahamas (Dix and Mullins, 1988a,b, 1992; Eberli, 1988; Freeman-Lynde et al., 1988; McClain et al., 1988). The spatial link between the toe-of-slope to deeper slope sediments on one side, and the upper slope to platform top deposits on the other side was closed in 1990 by the cores Unda and Clino of the Bahamas Drilling Project (Ginsburg, 2001a). The cores 1003 to 1007 of ODP Leg 166 completed the Bahamas transect along a single line from the platform top to the basin (Fig. 1). The relatively young age of the sediments cored, covering the Recent to Miocene, allows for examining early diagenetic features in all diagenetic zones with little later diagenetic overprint. Here we present a synopsis of diagenetic studies of the Bahamas transect cores conducted during the past 10 years. During this time, several thousand thin sections from Unda, Clino, 1003, 1005 and 1007 have been examined, the carbonate mineralogy of several thousand samples was determined with the XRD, carbon and oxygen stable isotopes have been measured for a similar number of samples, and several hundred samples have been investigated with scanning electron microscopy (SEM). Our investigations have covered large parts of the transect with detailed studies focused on the Miocene of the ODP cores and the entire recovered intervals of the BDP cores. The present synopsis is based on published articles (Melim et al., 1995, 2001a,b; Melim, 1996; Melim and Masaferro, 1997; Munnecke et al., 1997; Westphal and Munnecke, 1997; Westphal, 1998; Swart, 2000; Swart and Melim, 2000; Westphal et al., 1999b, 2000) and new, yet unpublished studies. In the rst part of this paper we will describe and characterize the diagenetic zones of the slope of Great Bahama Bank and describe the di erent styles of marine-burial diagenesis found along the

3 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 29 Fig. 1. Site map of the Bahamas transect showing Bahamas Drilling Project drill sites Unda and Clino and Ocean Drilling Program Leg 166 sites 1003 to Location of seismic line (Western Geophysical Line and westward extending seismic line, see Fig. 3), along which the drill sites are located is also shown. (From Eberli et al., 1997a.) transect. In the second part, we will address the paradigms that are questioned by these new ndings. 2. Lithofacies The Bahamas transect drilled the Neogene platform to slope sediments of the leeward side of Great Bahama Bank (Fig. 1; Eberli et al., 1997b; Ginsburg, 2001a,b). The transect extends from the present day platform top westward into the adjacent Santaren Channel (30 km from the present day platform margin) where cores were drilled in water depths of up to 660 m (Fig. 2). The lithofacies drilled along the Bahamas transect represent a variety of facies from shallow, reefdominated to deep water hemipelagic sediments

4 30 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 rich in planktic foraminifers. This summary of lithofacies is largely derived from published work (see Beach and Ginsburg, 1980; Eberli et al., 1997b; Eberli, 2000; Betzler et al., 1999, 2000; Westphal, 1998; Kenter et al., 2001; and Manfrino and Ginsburg, 2001). The upper portion of Great Bahama Bank is composed of shallow-water ramp, platform and reefal facies. Great Bahama Bank is currently a at-topped platform but seismic and core studies revealed an older ramp pro le that evolved into a at-topped platform during the Pliocene (Beach and Ginsburg, 1980; Schlager and Ginsburg, 1981; Beach, 1982; McNeill et al., 1988; Eberli and Ginsburg, 1987, 1989). Skeletal packstone to grainstone typi es the ramp facies whilst the platform facies are characterized by shallowing-upward packages of peloidal to skeletal wackestone to grainstone and/or coral framestone (Beach and Ginsburg, 1980; Kenter et al., 2001; Manfrino and Ginsburg, 2001). Subaerial exposure horizons are common, particularly in platform facies (Beach and Ginsburg, 1980; Beach, 1995; Manfrino and Ginsburg, 2001). The deeper forereef and deeper margin facies forms a transition from the bank top to the upper slope. This facies was recovered in cores Clino and Unda and is characterized by platform-derived ne-grained skeletal to mixed skeletal and non-skeletal wackestone to grainstone that alternate with coarse-grained intervals. Both ne- and coarse-sand intervals are very similar in grain composition (60% non-skeletal and 40% skeletal grains; Kenter et al., 2001). Unlike the shallowwater facies, these deeper water deposits do not contain subaerial exposure horizons. Instead phosphatic hardgrounds and rmgrounds punctuate the succession (Kenter et al., 2001; Melim et al., 2001b). The majority of the Bahamas transect is in slope to basin facies. In the Lower Pliocene^Miocene there are three main lithofacies found: (1) light-gray wackestones to packstones characterized by shallow-water bioclasts; (2) dark-gray wackestones characterized by increased pelagic components; and (3) grainstones to packstones with shallow-water bioclasts interpreted as turbidites (Eberli et al., 1997a; Betzler et al., 1999; Kenter et al., 2001; Westphal et al., 1999a). The light-gray wackestones, with their shallow-water composition, are interpreted as highstand shedding when the carbonate factory of Great Bahama Bank was ooded and active (Betzler et al., 1999; Eberli, 2000; Kenter et al., 2001). The darkgray wackestones form during either lowstand or as condensed intervals during transgression (Betzler et al., 1999; Kenter et al., 2001). Turbidites form during all sea level positions (Betzler et al., 1999; Bernet et al., 2000; Eberli, 2000) with greater amounts during highstands (Bernet et al., 2000). When the western margin of Great Bahama Bank changed from a ramp pro le in the Lower Pliocene^Miocene to more of a platform bank in the Upper Pliocene^Pleistocene, the composition of the sediments on the bank top changed to peloidal (Beach and Ginsburg, 1980; Beach, 1982) which led, in turn, to more peloidal sediments on the slope (Westphal, 1998; Rendle et al., 2000; Kenter et al., 2001). The slope facies is composed of highstand deposits of monotonous ne-sand to silt-sized skeletal and peloidal grains, interrupted by intervals of coarse-grained skeletal sands interpreted as lowstand deposits (Eberli et al., 1997b; Kenter et al., 2001; Westphal, 1998; Head and Westphal, 1999). The aragonite-rich intervals in cores Clino and Unda and ODP sites 1003, 1004, 1005 and 1007 all begin near the base of the Upper Pliocene seismic sequence d (Fig. 2; Eberli, 2000; Eberli et al., 2001; Kenter et al., 2001). In Clino and Unda, margin progradation placed skeletal reef to platform facies above the peloidal interval whilst the more distal ODP sites remained peloidal to the present (Eberli et al., 1997a, 2001; Eberli, 2000; Kenter et al., 2001; Manfrino and Ginsburg, 2001), with important consequences for diagenetic potential (Melim et al., 1995; Rendle et al., 2000). 3. Geochemical data 3.1. Mineralogy The mineralogy data are presented in Figs. 2 and 3. The data for cores Clino and Unda are

5 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 31^34 Fig. 2. Mineralogy, seismic sequences, and diagenetic zones of the Bahamas transect superimposed on Western Geophysical Line and seismic line westward (location of seismic line is shown in Fig. 1). Compiled from Eberli et al., 1997a,b; Melim et al., 1995, 2002b; Kramer et al., 2000 and this study; with additional XRD data from T. Frank.

6 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 35 Fig. 3. Carbonate mineralogy and stable oxygen and carbon isotopes in BDP cores Unda and Clino. Note parallel trend from negative to positive oxygen and carbon isotopes. This shift is interpreted as the transition from meteoric to marine-burial diagenesis. (After Melim et al., 1995.) from Melim et al. (1995); the ODP Leg 166 data are from site chapters in Eberli et al. (1997a) augmented by unpublished data from P. Swart and T. Frank. The upper shallow-water facies in Clino and Unda are mainly low-mg calcite (LMC) with minor aragonite near the top of both cores (Fig. 3). The Miocene reef is extensively dolomitized (Fig. 3). The deeper water facies in all cores are characterized by LMC with minor aragonite and/or dolomite except for an aragonite-rich interval in Clino at W220^360 mbmp and the upper 100^150 m of ODP Leg 166 cores (Fig. 2). Site 1006 has greater aragonite at depth than the other cores (Fig. 2) Stable isotopes The bulk rock stable isotopic data for core Clino and Unda are presented in Fig. 3. There are three distinct intervals: (1) the upper portion of both cores (Clino 0^110 m; Unda 0^80 m) with negative carbon and oxygen isotopic compositions; (2) a transition interval where the isotopic values progressively shift downcore toward positive isotopic compositions (Clino 110^145; Unda 80^130 m); and (3) the rest of both cores with positive compositions (Fig. 3; Melim et al., 1995, 2001b; Melim and Masaferro, 1997). Dolomiterich intervals have more positive oxygen isotopic

7 36 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 As might be expected from such a wide range of depositional environments, the sediments of the Bahamas transect have altered in a variety of diagenetic environments. We have identi ed meteoric, mixing-zone, and phreatic-marine diagenesis. In addition, the phreatic-marine environment can be divided into sea oor diagenesis, marineburial diagenesis, and deep burial diagenesis. We describe each of these before discussing the implications of these results for existing carbonate diagenetic paradigms Meteoric diagenesis Fig. 4. Schematic of contrasting styles of marine-burial diagenesis of a skeletal grainstone depending on the permeability of the surrounding sediment. (A) Starting sediment. (B) Open-system marine-burial diagenesis in intervals with high permeability. Aragonite is dissolved and is removed from the system, leaving a highly porous, poorly cemented limestone consisting predominantly of low-mg calcite. (C) Closed-system marine-burial diagenesis in intervals with lower permeability. Aragonite grains are either dissolved or replaced by neomorphic spar. Blocky calcite spar occludes most porosity. (After Melim et al., 1995.) compositions (up to 4x; Melim et al., 2001b) as dolomite in these cores in enriched approximately 3 permil relative to calcite (Swart and Melim, 2000). 4. Diagenesis along the Bahamas transect 4.1. Introduction The meteoric diagenetic zone occurs in the upper 100^150 m of cores Clino and Unda. It is characterized by complete alteration of an aragonite-rich original sediment to a low-mg calcitic limestone (Figs. 3 and 4; Melim and Masaferro, 1997; Melim et al., 2001b). Bulk rock isotopic data show depleted carbon and oxygen values average N 18 O=33.0 þ 0.7x; N 13 C= 31.6 þ 1.7x), typical of meteoric diagenesis (Fig. 3; Melim et al., 2001b). Grainstones to packstones typically have neomorphism, micrite envelopes, and blocky calcite spar cements (Plate IA). Minor meniscus cements are also found. Finer grained wackestones to mudstones have been altered to dense micrite and microspar, often with moldic porosity. Laminated crusts, root casts, circumgranular cracking and blackened grains document subaerial exposure surfaces (caliches) (Manfrino and Ginsburg, 2001; Melim et al., 2001b) Mixing-zone diagenesis Directly underlying the meteoric diagenetic zone, there is a 40^50 m transition interval where the oxygen isotopic values gradually shift from negative values characteristic of meteoric diagenesis to positive values (average N 18 O = +0.9x) indicative of marine phreatic diagenesis (Fig. 3; Melim et al., 1995, 2001b). The top of this interval occurs at di erent depths in Clino and Unda but in both cases begins approximately 10 m below the deepest subaerial exposure surface (Fig. 3; Melim and Masaferro, 1997). We interpret this transition interval as forming in a marine^meteoric mixing-zone during development of the overlying subaerial exposure surface.

8 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 37 Plate I. Photomicrographs of thin sections from the Bahama transect cores. (A) Shallow-water grainstones altered to low-mg calcite with blocky spar, moldic porosity (mostly lled) and micrite envelopes. This sample has negative oxygen isotopic values indicating diagenesis in meteoric pore uids. Location Unda, 58.8 mbmp. (B) Upper slope grainstones to packstones with extensive moldic porosity and micrite envelopes but only minor dogtooth and overgrowth cementation. This sample has positive oxygen isotopic values, indicating diagenesis in marine pore uids. Location Unda, mbmp. (C) Upper slope peloidal grainstone with microspar matrix, neomorphism, blocky spar cementation and moldic porosity. This sample has positive oxygen isotopic values, indicating diagenesis in marine pore uids. Location Clino, mbmp. (D) Skeletal grainstone altered in the open-system style of marine-burial diagenesis. Aragonitic skeletal grains are dissolved forming moldic porosity, with or without micrite envelopes. Minimal cementation. Interval A-36X-1, 20^22 cm; depth m. (E) Skeletal grainstone altered in the closedsystem style of marine-burial diagenesis. Aragonitic skeletal grains are often neomorphosed to pale yellow blocky calcite and primary and secondary pores are nearly complete occluded by blocky calcite spar. The result has very low permeability. Interval B-56X-1, 28^32 cm; depth m.

9 38 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 In core Clino, the mixing-zone also displays a change in mineralogy (Fig. 3). The meteoric zone is entirely low-mg calcite as is much of the mixing-zone interval. Near the base of the mixing zone, 5^10% aragonite occurs (as N 18 O reaches 0x) and 2^5% dolomite marks the top of the underlying marine diagenetic zone (N 18 O= +1x; Fig. 3). In core Unda, the mixing-zone overlaps a rmground with penecontemporaneous dolomite that predates the deposition of the overlying shallow-water facies (Melim et al., 1995; Swart and Melim, 2000). This earlier dolomite obscures the mineralogic and isotopic changes through the mixing-zone interval (Fig. 3). However, as in Clino, the platform interval is entirely low-mg calcite with negative stable isotopic values and the underlying upper slope facies is low-mg calcite with minor amounts of dolomite and aragonite and positive stable isotopic values. Petrographically, the mixing-zone interval is characterized by extensive moldic porosity, blocky or dogtooth cements, and micrite or microspar. Cementation is relatively minor compared with the amount of dissolution present Marine phreatic diagenesis Sea oor diagenesis Rare examples of isopachous cement occur in the platform facies that probably formed during sea oor diagenesis (Melim et al., 2001b). More signi cant and widespread sea oor diagenesis occurs as marine hardgrounds found in all cores (Eberli et al., 1997a,b; Melim et al., 2001b). Evidence for sea- oor lithi cation includes phosphatized and blackened surfaces, borings, and reworked pebbles in overlying units (Eberli et al., 1997a,b; Melim et al., 2001b) Marine-burial diagenesis The majority of the Bahamas transect recovered upper slope to lower slope facies, almost all of which were exclusively altered in marine pore uids (Fig. 2). This diagenesis in marine pore uids mimics many aspects of diagenesis in meteoric pore uids, most notably by producing a mainly low-mg calcite limestone with blocky spar, neomorphism, microspar and moldic porosity. We term this diagenesis marine-burial diagenesis (Melim et al., 1995) to distinguish it from both the well-documented near-surface marine diagenesis characterized by hardgrounds and/or marine cementation (e.g. James and Choquette, 1990a) and deeper burial diagenesis characterized by compaction, pressure solution, and late cements (e.g. Scholle and Halley, 1985; Choquette and James, 1990). Whilst the meteoric diagenetic environment has negative stable isotopic values (Fig. 3), the marine-burial environment in cores Clino and Unda has positive stable isotopic values (Fig. 3; average N 18 O = +0.9 þ 0.3x; N 13 C = +3.0 þ 0.9x; Melim et al., 1995, 2001b; Melim and Masaferro, 1997). Oxygen isotopic values are mainly a function of water composition and temperature (Anderson and Arthur, 1983). Given the relatively short core distance over which the transition occurs, temperature alone cannot account for the approximate 4x shift in N 18 O (Melim et al., Plate II. SEM micrographs of samples from the Bahamas transect cores. All samples are polished and slightly etched prior to gold coating. (A) Upper slope sample with tight mosaic of microspar cement with engulfed aragonite needles. Location Clino, mbmp. (B) Undeformed dino agellate cyst in cemented limestone from the upper slope. Spherical preservation of the cyst implies early lithi cation. Clino, mbmp. (C) Uncemented layer from the upper slope consisting largely of platform-derived aragonite needles. Note poor preservation of aragonite needles that points to partial dissolution of these metastable constituents (sample is not etched!). Location Clino, mbmp. (D) Turbidite from the lower slope of the Bahamas transect that shows aragonite needles enclosed in tight microspar cement. This diagenetic style is reminiscent of cemented samples from the upper slope. Interval C-34R-3, 125^129 cm; depth m. (E) Sample from dark layer shows foraminifer tests that collapsed due to mechanical compaction. Interval C-21R-5, 03^08 cm; depth m. (F) Fine-grained, uncemented matrix of dark, compacted layer. Note presence of coccoliths. Interval C-23R-2, 19^23 cm; depth m. (G) Detail of coccolith-rich matrix of dark layer with high micro porosity. Interval C-21R-5, 3^8 cm; depth m. (H) Light layer with microsparitic cementation and moldic porosity. Interval C-2R-1, 65^69 cm; depth m.

10 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 39

11 40 L.A. Melim et al. / Marine Geology 185 (2002) 27^ b). Therefore, the change must re ect changes in the N 18 O composition of the water from a meteoric to marine composition (Melim et al., 1995, 2001b). Frank (2000) and Frank and Bernet (2000), also found positive N 18 O values for Miocene samples in sites 1006 and 1007 and interpret them to support alteration in seawater during early burial. There are two basic styles of marine-burial diagenesis that appear to be controlled by the permeability of the surrounding sediments (Melim et al., 1995, 2001b; Melim and Masaferro, 1997). Although we de ne two distinct styles herein (Fig. 4), it is important to recognize that a complete gradation exists between the two end members. In addition, up to 100% dolomite is present, particularly in core Unda (Fig. 2). For this paper, we will restrict ourselves to the calcium carbonate portions; see Swart and Melim (2000) for a discussion of the dolomite High-permeability intervals Skeletal grainstones to packstones start with relatively high porosity and permeability (Enos and Sawatsky, 1981). This allows easy movement of pore uids and results in an open-system style of marine-burial diagenesis characterized by extensive secondary porosity and minimal cementation (Melim et al., 1995, 2001a). Aragonitic skeletal grains are dissolved forming moldic porosity, with or without micrite envelopes (Plate IB,D). The micrite rims probably formed prior to deposition in the deeper water environment since they are much less common than in shallow-water facies. Nevertheless, some micritization in the slope environment cannot be excluded. Aragonitic peloids are either dissolved or preserved, forming the 5^10% aragonite common in these intervals (Fig. 2). Aragonitic peloids resist dissolution more than do aragonitic skeletal grains. The reason is unknown, but perhaps organic coatings of some kind isolate the peloids from the pore uids. Cementation is limited to minor dogtooth to ne blocky calcite spar and traces of overgrowth cements (Plate IB). As a result, cementation is minimal and lithi cation is poor. This style of alteration is found in all of the Bahamas transect cores except site 1006 but is most common in Unda because the sediments in this proximal core are coarser grained (Kenter et al., 2001). In the core 1003 and 1005, the seismically transparent interval that coincides with seismic sequence f is very poorly lithi ed (Shipboard Scienti c Party, 1997a,b; Anselmetti et al., 2000) but the near absence of aragonite (Fig. 2) attests to extensive alteration. On closer examination, moldic porosity is abundant showing near complete dissolution of aragonitic (and presumably high-mg calcite) skeletal grains without signi cant cementation (Plate ID). The high secondary porosity without cementation requires wholesale exportation of aragonite out of the system (hence, open-system diagenesis), presumably into the ocean. Some of the calcium carbonate might be taken up by cementation of ne-grained beds, but the intervals altered with this open-system style of marine-burial diagenesis are not associated with su cient cemented beds to account for the amount of aragonite dissolved. This contradicts, at least for this interval, the suggestion of Kramer et al. (2000) that the pore uids in Leg 166 are in situ, as substantial ow is required to remove the dissolved components Low-permeability intervals Most of the slope facies along the transect are ne-grained packstones to wackestones with interbedded turbidite grainstones (Eberli et al., 1997a; Betzler et al., 1999; Kenter et al., 2001). The overall ner grain size, as compared to the open-system intervals described above, produces lower permeability (Melim et al., 2001a) and a di erent style of marine-burial diagenesis characterized by a more closed-system recycling of calcium carbonate. The variable lithology in the slope facies produces variable styles of marine-burial diagenesis within the context of a generally closed system. The greatest di erence is between a peloid-dominated interval (Upper Pliocene seismic sequence d in Clino and Unda; Upper Pliocene to Recent seismic sequences a, b, c and d in 1003, 1005 and 1007; Shipboard Scienti c Party, 1997a,b,c; Eberli et al., 1997b; Eberli, 2000) and more skeletal to peloidal intervals deeper in the cores (Shipboard Scienti c Party, 1997a,b,c; Eberli et al.,

12 L.A. Melim et al. / Marine Geology 185 (2002) 27^ b; Eberli, 2000; Kenter et al., 2001). In the peloid-rich interval, the easily deformed soft peloids compacted very early resulting in relatively high microporosity (average 36% in Clino) but very low permeability (average 3 md in Clino) (Melim et al., 2001a). This very low permeability slowed, or perhaps even stopped, uid movement and partially protected the sediment from alteration. As a result, the peloid-rich interval has a high aragonite content (Shipboard Scienti c Party, 1997a,b,c; Rendle et al., 2000), even where it is deeper in Clino (Figs. 2 and 3; Melim et al., 1995, 2001b; Eberli et al., 1997b). This aragonite consists of well-preserved aragonite needles that the bulk of which are thought to have composed the peloids (Westphal, 1998; de Mol et al., 1998; Rendle et al., 2000). The light-gray, shallow-water derived, skeletal to peloidal packstones to wackestones that comprise the majority of the ODP cores and the deeper portion (below 368 m) of Clino also altered with a closed-system style of marine-burial diagenesis. Aragonitic skeletal grains are either dissolved or altered to neomorphic calcite spar. The molds are usually partially to completely lled with ne blocky calcite spar. Pelagic foraminifera commonly show overgrowths. Celestite is common in trace amounts as either a cement or replacement of aragonite. Compaction in the lightgray beds is highly variable. Although Betzler et al. (1999) describe them as uncompacted in the ODP sites, palynomorphs in 1003 and 1007 show minor compaction. In Clino, these beds vary from uncompacted (Plate IIB) to extensively compacted (Westphal et al., 2000). This variability of the light-gray beds is in contrast to the darkgray beds that always are strongly compacted (Betzler et al., 1999). The matrix of the light-gray packstones to wackestones and the more altered portions of the peloidal interval is calcite microspar. Especially in samples from the upper slope sites (Munnecke et al., 1997; Westphal, 1998), but also less commonly from the deeper settings (Kenter et al., 2002), SEM examinations revealed that the microspar encloses aragonite needles (Plate IIA,D) and exhibits sharp boundaries towards larger components. These features indicate that the microspar found in these samples is a cement (Munnecke et al., 1997). The more compacted, dark-gray pelagic-derived packstones to wackestones do not contain molds and have a coccolith-rich matrix without aragonite needles (unlike the light-gray beds that contain both). Despite the lack of aragonite needles, the dark-gray beds contain higher amounts of aragonite than surrounding light-gray beds. The dark-gray beds also have higher clay content (Betzler et al., 1999). Compaction is seen in broken pelagic foraminifera, squished infaunal pellets, deformed burrows and palynomorphs (Betzler et al., 1999; Melim et al., 2001b; and own new observations). Cement of any kind is very rare. Based on their coarse grain size and lack of matrix, the interbedded turbidite grainstones probably started with high permeability (e.g. Enos and Sawatsky, 1981). They do not, however, show the same open-system style of diagenesis that characterizes the generally higher permeability intervals (Melim et al., 1995, 2001b). Rather, they contain extensive blocky spar cementation and aragonite neomorphism (Plates IE and IIA) similar to the cemented layers in the low-permeability intervals in upper slope settings (Clino). The degree of cementation in these beds requires some local importation of calcium carbonate as almost all primary and many secondary pores are lled. In the case of the grainstones interbedded with aragonite-rich peloidal interval, adjacent peloid beds are more altered than those farther away indicating they acted as donor beds. Presumably nearby beds acted in a similar fashion in the more skeletal-peloidal intervals Deep-burial diagenesis The transition from marine-burial diagenesis, a relatively near-surface process, and deep-burial diagenesis is very di cult to pinpoint. However, compaction gradually increases with depth in the ner-grained intervals. For example, cracked and broken pelagic foraminifera are common in the dark-gray beds in the Miocene of Clino and Leg 166 cores but absent in the Upper Pliocene (Eberli et al., 1997a; Betzler et al., 1999; Kenter et al., 2001; this study). In addition, pressure solution seams are present below W1000 m in sites 1003

13 42 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 and 1007 (Shipboard Scienti c Party, 1997a,b). Below 300^350 m, minor celestite- lled fractures were found in lithi ed intervals of Clino, Unda (Melim et al., 2001b) and sites 1003, 1005, and 1007 (but not 1006) (core descriptions in Eberli et al., 1997a). 5. Questioning paradigms 5.1. Limitations on lowstandmeteoric diagenesis The upper portion of Great Bahama Bank was extensively altered by meteoric diagenesis during Pleistocene and older sea-level lowstands (e.g. Williams, 1985; Beach, 1995). Considering the large-scale sea-level falls of the Pleistocene (down to 120 m; Fairbanks, 1989), the potential exists for meteoric alteration well below the platform top (Beach, 1995). Earlier deep core borings found meteoric alteration to 5000^6000 m (Spencer, 1967; Meyerho and Hatten, 1974; Walles, 1993). However, these deep cores were drilled on the platform top where shallow-water deposition with periodic emergence has been the rule since the Cretaceous (Spencer, 1967; Meyerho and Hatten, 1974; Walles, 1993), thus making it impossible to determine burial depth during meteoric alteration (Melim and Masaferro, 1997). The Bahamas transect cores, on the other hand, extend beneath the shallow-water facies into underlying upper to lower slope facies deposited below possible subaerial exposure (see Eberli et al., 1997a,b; Ginsburg, 2001a). The two proximal cores recovered the transition from shallow-water platform facies into upper slope facies, thus allowing a test of the models for maximum burial depth of meteoric diagenesis. The lower limit of meteoric in uence can be de ned in Clino and Unda based on the stable isotopic data. If we take the most generous de nition, the base of the mixing-zone, then meteoricin uenced diagenesis extends to 145 m in Clino and W130 m in Unda (Fig. 3; Melim, 1996). If we take a more literal de nition, the top of the mixing-zone, then meteoric diagenesis only extends down to W110 m in Clino and W80 m in Unda (Fig. 3; Melim, 1996). Melim (1996) and Melim and Masaferro (1997) document a similar mixing-zone between 32 m and 77 m in a core from Key West, Florida (Florida Geological Survey core Stock Island). In addition, the top of the mixing-zone occurs W10 m below the deepest subaerial exposure surface recovered in all three cores. The latest Pleistocene sea-level lowstand was 3120 m (Fairbanks, 1989). The water table fairly closely tracks sea level in carbonate islands (Vacher, 1988). Therefore, if the latest Pleistocene sealevel lowstand had a chemically active meteoric lens, we should nd meteoric diagenesis at and below this depth. Instead, 3120 m is near the top of the mixing zone in Clino, near the base of the mixing zone in Unda, and over 40 m below the base of the mixing zone in Stock Island (Melim, 1996; Melim and Masaferro, 1997). Apparently, the latest Pleistocene sea-level lowstand produced no phreatic meteoric diagenesis. There is aragonite present at this depth in Unda (Fig. 3; also in Stock Island; Melim, 1996; McNeill et al., 1996; Melim and Masaferro, 1997), so the diagenetic potential is still there. We cannot say there was not a phreatic lens, only that any lens present left no diagenetic signature behind. In addition, there is no diagenetic record from any sea-level lowstand that reached to similar depths (below 32^77 m in Stock Island). Melim (1996) proposed two factors that could lead to a diagenetically inactive lens at greater depths: (1) the greater percolation distance allows the water to reach saturation prior to entering the lens, and (2) the large distance exceeds the reach of soil-derived organic matter, known to drive diagenesis within meteoric lenses (Smart et al., 1988; McClain et al., 1992). If large-scale sea-level lowstands are not responsible for major phreatic diagenesis, the observed alteration must occur at other times, presumably when the meteoric lens occurs closer to the surface. Since the top of the mixing zone occurs 10 m below the rst subaerial exposure surface, we interpret the deepest meteoric diagenesis to the rst exposure of each core location, not to later, perhaps larger, sea-level lowstands. In these cores, the limit of pure meteoric diagenesis is W10 m below the land surface. This is consistent with modern hydrogeochemical studies in the

14 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 43 Bahamas (Halley and Harris, 1979; McClain et al., 1992) and with the thickness of many modern Bahamian meteoric lenses (Vacher and Wallis, 1992; Whitaker and Smart, 1997). McClain et al. (1992) cautioned against using distribution of meteoric fabrics to interpret paleophreatic lens distribution as their results showed alteration only in the upper portion of the lens. Our results extend that caution as larger sea-level lowstands may not be recorded at all (Melim, 1996) Characteristics of mixing-zone diagenesis Current models of mixing-zone diagenesis are based on studies from coastal regions near groundwater discharge points (Hanshaw and Back, 1980; Smart et al., 1988; Ward and Halley, 1985) or from islands with small, areally restricted (1^10 km diameter) phreatic lenses (Budd, 1988; Anthony et al., 1989). Also of interest, however, is the behavior of mixing-zones developed during sea-level lowstands when entire carbonate platforms are exposed. Various workers have proposed that patterns recognized on islands and in coastal regions can be projected through time and space to predict diagenesis such as cavernous porosity and/or dolomitization in extended mixing zones (e.g. Badiozamani, 1973; Humphrey and Quinn, 1989; Beach, 1995). The results from Clino and Unda do not support either of the existing models for mixing-zone diagenesis. Whilst dissolution does occur, it is fabric-selective dissolution of aragonite grains essentially identical to that found in the underlying marine-burial diagenetic zone. The only exception is that mixing-zone diagenesis is more e cient at removing the last few percent of aragonite that marine-burial diagenesis usually leaves untouched (Fig. 3). Cavernous to vuggy porosity does not occur within the mixing zone, although it is common in the overlying meteoric diagenetic zone (Melim et al., 2001a,b). Dolomite, on the other hand, is completely absent from the mixing zone (Fig. 3). The earlier sea oor dolomitization in Unda obscures this relationship. In Clino, lacking this earlier dolomite, the rst occurrence of trace amounts of dolomite downcore is an excellent proxy for the rst occurrence of marine isotopic values (N 18 O=+1x). In other words, dolomite occurs within the marineburial diagenetic zone but not within the mixing zone. The coastal-mixing zone is likely not a good analog for the deeper-mixing zone for two reasons: (1) uid ux is much higher through the coastal-mixing zone because of discharging groundwater (Sanford and Konikow, 1989); (2) the proximity to the land surface allows input of organic matter from the surface, shown by Smart et al. (1988) enhance dissolution. In addition, the coastal environment is a much more hydrochemically complex region where groundwater discharging from the meteoric lens mixes not only with seawater but also with freshwater in ltrating from the surface. Given these signi cant di erences, it should come as no surprise that the extended mixing-zone environment shows di erent diagenesis Marine-burial diagenesis mimics meteoric diagenesis Petrographic fabrics are frequently used for determining the diagenetic environment of carbonate rocks, although most workers recognize the need for additional geochemical lines of evidence. Features that are characteristic of the phreatic meteoric environment include aragonite dissolution (molds), neomorphism, blocky spar (also known to form during burial), microspar, and isopachous equant cements composed of low-mg calcite (Folk and Land, 1975; Flu«gel, 1982; Tucker and Wright, 1990; James and Choquette, 1990b). Several workers have identi ed aragonite dissolution accompanied by blocky spar cementation in deep marine sediments, and attributed it to deep, cold undersaturated waters (i.e. below the aragonite compensation depth; Halley et al., 1984; Freeman-Lynde et al., 1988; Dix and Mullins, 1988a,b; Frank and Bernet, 2000) or to deep meteoric ow (Enos, 1988). Workers in shallowwater environments, therefore, have continued to assume that these fabrics equal meteoric diagenesis; in some cases even when the isotopic evidence was equivocal (e.g. Mutti, 1995). Results from cores along the Bahamas transect

15 44 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 show that all these petrographic patterns also can and commonly do form in marine pore uids, at depths well above the aragonite compensation depth. Petrographic characteristics display patterns identical to those long thought to identify phreatic meteoric diagenesis. Neomorphic Halimeda, aragonite leaching (moldic porosity), micritic envelopes, blocky spar, dogtooth spar, and microspar cements have been observed along the transect down to the toe of slope (Melim et al., 1995, 2001b; Westphal, 1998, and own new observations). Stable isotope analyses showed, however, that these features formed in marine pore uids and were never in uenced by meteoric uids (Fig. 3; Melim et al., 1995). In addition, the distance from the platform margin makes it unlikely that meteoric water could reach the deeper slope sites; our data indicates it did not. Limestones in the Clino and Unda completely altered by marine pore uids occur within 100^150 m of sea level (and 75 m in the Stock Island Core, Florida); well above not only the modern aragonite compensation depth ( m; Droxler et al., 1988) but any possible ancient aragonite compensation depth. The diagenetic environment responsible for this alteration has been termed the marine-burial environment (Melim et al., 1995, 2001b). Marineburial diagenesis occurs after sea oor diagenesis and before deeper burial processes such as pressure-solution. Since seawater above the aragonite compensation depth is saturated with respect to aragonite, some mechanism must be identi ed to drive marine-burial diagenesis. We suggest the chemical environment responsible for dissolution and reprecipitation in the marine-burial realm is triggered by microbially induced decay of organic matter. Elevated CO 2 levels initiate aragonite dissolution and reprecipitation of the calcium carbonate as calcite cement, resulting in early chemical and mechanical stabilization of metastable carbonates. This interpretation disagrees with that of Frank and Bernet (2000) who suggest the aragonite compensation depth was shallow enough in the Miocene to allow undersaturated waters into 1007 and 1006 (despite this requiring a change of more than 3000 m). Considering the similarity of 1007 to substantially shallower cores (e.g. Clino), we reject a model that cannot explain all of the marine-burial diagenesis observed. Marine-burial diagenesis is intermediate between sea oor diagenesis (with essentially unaltered seawater; James and Choquette, 1990a; Tucker and Wright, 1990) and deep burial diagenesis (typically including compaction and alteration with evolved pore uids; Scholle and Halley, 1985; Choquette and James, 1990; Tucker and Wright, 1990). In order for marine-burial diagenesis to start, saturated seawater must be driven to undersaturation with respect to aragonite. Once this occurs, the di erent solubility of aragonite and high-mg calcite compared to low-mg calcite could sustain the reaction going until all of the aragonite is consumed (Budd, 1988; James and Choquette, 1990a). Thus, the reaction may start at very shallow sub-surface depths and continue into the deeper burial environment until the supply of metastable minerals such as aragonite and high-mg calcite becomes exhausted. Our data indicate just such a transition but leave open the question of where early starts. The new ndings that many fabrics are not unequivocal indicators of the meteoric environment should caution us not to entirely rely on petrographic evidence for interpretations of the fossil record but to base interpretations on additional evidence such as subaerial exposure surfaces or geochemical evidence. Although most of the intervals studied were upper to lower slope facies, the base of Unda (443.49^ m) is interpreted as deeper shelf to platform (Kenter et al., 2001) and was also altered during marine-burial diagenesis. So marine-burial diagenesis is probably not limited to deeper water facies. We speculate that many ancient successions have been interpreted with a bias towards meteoric diagenesis and need to be re-evaluated. The similarity of fresh water diagenesis and shallow marine-burial diagenesis is caused mostly by the fact that the dominant source for carbonate cement is the selective dissolution of sedimentary aragonite: Generally, the shape and mineralogical composition of calcium carbonate precipitates strongly depends on the Mg/Ca ratio of the precipitating uid (Folk, 1974; Given and Wilkinson, 1985). Fresh water has a low ratio

16 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 45 and thus the precipitates are composed of calcite characterized by a typical blocky shape. Seawater normally has a higher Mg/Ca ratio and thus the precipitates (aragonite and/or high-magnesian calcite) are typically elongated (along the c-axis) or needle-shaped (Folk, 1974; Given and Wilkinson, 1985). The microspar crystals and also the sparitic cement in the coarse grained limestones show a blocky fabric. Principally, four sources for the carbonate are possible: (a) fresh water ^ this can be excluded by the stable isotope data; (b) marine-derived pore water ^ such waters have a high Mg/Ca ratio and thus their precipitates should be elongated rather than blocky; (c) pressure solution ^ unrealistic in the shallow burial realm; (d) selective dissolution of aragonite ^ supported by the common association of moldic porosity. This dissolution would shift the pore water geochemistry towards low Mg/Ca ratios because aragonite is a very Mg-poor carbonate phase. The resulting pore water would be ^ to some extent ^ similar to fresh water and thus the precipitates (blocky spar, microspar) are very similar in shape, resulting in the above mentioned overestimation of fresh water diagenesis in the geological record Microspar as a cement Petrographic observations, that matrix microspar crystals enclose aragonite needles (Plate IIA,D) and exhibit sharp boundaries towards larger components, and that samples cemented by microspar are uncompacted, indicate that the microspar found in these samples is a cement (Munnecke et al., 1997; Westphal et al., 2000). These ndings are in contradiction to the widely accepted interpretation that microspar is the product of recrystallization (Folk, 1959, 1965). On the basis of light microscopic examinations, Folk (1959, 1965) suggested that microspar forms by recrystallization from a previously lithi ed micrite ( aggrading neomorphism ). In his model, Mg-ions, that are released into the pore water during early diagenetic alteration of high-mg calcite to low-mg calcite, form an Mg-cage around the micrite crystals and thereby inhibit growth of these crystals (Folk, 1959, 1965). When the Mgions are removed from the pore water, usually by freshwater in ltration, the micrite crystals start to grow until they reach microspar size. Folk s theory, however, fails to explain why, e.g., chalk remains unaltered when it is exposed to meteoric diagenesis (James and Choquette, 1983), and it o ers no explanation for the source of the calcium carbonate required for the lithi cation of the limestones. With SEM, Lasemi and Sandberg (1984) recognized in Pleistocene aragonite-dominated carbonate muds from the Bahamas and South Florida, that microspar can form as a cement during meteoric diagenesis. Calcite crystals with mean diameters between 5 and 15 Wm (microspar) are precipitated within the sediment, and small carbonate grains (e.g. aragonite needles) are engulfed in these microspar crystals. SEM examinations of the ne-grained slope deposits along the Bahamas transect revealed striking textural similarities to the Pleistocene samples of Lasemi and Sandberg (1984) (Munnecke et al., 1997; Westphal, 1998). Based mainly on the textural observations, it is thought that microspar cement formed by a process that is fundamentally di erent from the process of aggrading neomorphism proposed previously (Folk, 1959, 1965, 1974). The microspar crystals form a tightly tted mosaic engul ng undeformed components and aragonite needles of the matrix, indicating that they represent an early cement that has lithi ed the pristine aragonite-dominated carbonate mud (Fig. 5; Munnecke et al., 1997). Moreover, the process for microspar formation presented here explains why calcitic bioclasts show sharp boundaries with the microspar matrix. Aggrading neomorphism (recrystallization) fails to explain this texture. Additionally, recrystallization from low- Mg calcitic micrite to microspar is energetically highly improbable. Once a stable low-mg calcitic composition is reached, little driving force remains for recrystallization (Veizer, 1977; Steinen, 1978; Sandberg and Hudson, 1983). In contrast to the Pleistocene samples of Lasemi and Sandberg (1984), the Pliocene sequences from Clino were lithi ed in the marine-burial environment. The textural features observed suggest that microspar cement sensu Lasemi and Sandberg (1984) not only occurs in meteoric diagenetic

17 46 L.A. Melim et al. / Marine Geology 185 (2002) 27^53 Strontium is incorporated into the structure of both calcite and aragonite, substituting for Ca in the crystal lattice. Generally speaking aragonite has higher concentrations of Sr (7000^9000 ppm), while HMC and LMC (1000^4000 ppm) have lower values (Milliman, 1974). During the recrystallization of aragonite, HMC, and LMC to inorganic forms of LMC and dolomite, Sr is excluded from crystal structure. Hence in an open system the Sr concentration of the nal diagenetic product will be dictated by the Sr/Ca ratio in the solution and the distribution coe cient of the mineral in question (Veizer, 1983). The distribution coe cient is de ned as the ratio of the trace element to calcium in the mineral divided by the same ratio in the solution. D Sr ¼ Sr=Ca ðmineralþ Sr=Ca ðfluidþ Fig. 5. Reconstruction of microspar cement development in aragonite-dominated carbonate mud. (A) Unlithi ed aragonite needle mud. (B) Beginning of precipitation of microspar. (C) Completely cemented sediment. (D) Empty pits in microspar resulting from dissolution of aragonite needles. (E) Mature microsparitic limestone (after Westphal, 1998; Munnecke et al., 1997; Munnecke and Samtleben, 1996). environments, but also can be formed during marine-burial diagenesis. Thus, similar to other cement types and to many diagenetic features, the environmental signi cance of microspar is limited (Munnecke and Samtleben, 1996; Munnecke et al., 1997; Westphal, 1998) Strontium content as an indicator of diagenesis In the case of seawater the D Sr for aragonite is approximately unity, meaning that there is no active accumulation or discrimination of the element into this phase. In contrast the D Sr for organic LMC is 0.12 and for inorganic LMC is 0.05 (Kinsman, 1969; Veizer, 1983). The D Sr for dolomite is also about (Vahrenkamp and Swart, 1994) and is dependent upon the stoichiometry of the dolomite. In the past, low Sr concentrations in carbonates have been taken as an indicator of alteration in a freshwater regime (Land and Epstein, 1970; Gross, 1964). High Sr concentrations have been suggested in cements produced in hypersaline environments (Land and Hoops, 1973; Veizer et al., 1978). Although such generalities are not completely without merit, data on the chemical composition of pore waters from ODP and DSDP show that marine waters can show a wide variety of Sr/Ca ratios, which would ultimately, result in a range of Sr concentrations in diagenetic carbonates formed in these environments. In an open system, the Sr concentration of LMC formed at 25 C from a solution of actively circulating seawater is about 500 ppm. In deep sea pelagic sediments which overlay basalts, alteration of the igneous rocks produces pore uids with very high Ca 2þ concentrations (50^60 mm) resulting in uids with low Sr 2þ /Ca 2þ which in turn result in the precipitation of cements with low Sr concentrations (Baker et al., 1982). In closed systems una ected by basalt alteration, Sr excluded during the recrystallization process can build up to high concentrations resulting in diagenetic calcites and dolomites with very high val-