Questioning carbonate diagenetic paradigms

Transcription

Marine Geology 185 (2002) 27^53
www.elsevier.com/locate/margeo
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
a
Western Illinois University, Macomb, IL 61455, USA
Institut fu«r Geologie und Pala«ontologie, Universita«t Hannover, Callinstr. 30, 30167 Hannover, Germany
c
RSMAS, University of Miami, 4600 Rickenbacker Cswy., Miami, FL 33149, USA
Institut fu«r Geologie und Pala«ontologie, Universita«t Tu«bingen, Herrenberger Str. 51, 72070 Tu«bingen, Germany
b
d
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 N18 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.
E-mail address: la-melim@wiu.edu (L.A. Melim).
0025-3227 / 02 / $ ^ see front matter B 2002 Elsevier Science B.V. All rights reserved.
PII: S 0 0 2 5 - 3 2 2 7 ( 0 1 ) 0 0 2 8 9 - 4
MARGO 3048 11-6-02 Cyaan Magenta Geel Zwart
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. B 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
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 1007. 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
30
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
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.
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).
3.2. 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
36
genetic 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.
4.2. 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
As might be expected from such a wide range
of depositional environments, the sediments of the
Bahamas transect have altered in a variety of dia-
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 N18 O = 33.0 R 0.7x; N13 C =
31.6 R 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).
4.3. 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 N18 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.
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, 443.53 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, 225.56 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 166-1005A-36X-1, 20^22 cm; depth 296.3 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
166-1003B-56X-1, 28^32 cm; depth 522.68 m.
38
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 N18 O reaches
0x) and 2^5% dolomite marks the top of
the underlying marine diagenetic zone (N18 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.
4.4. Marine phreatic diagenesis
4.4.1. 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). Evi-
dence 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).
4.4.2. 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 N18 O = +0.9 R 0.3x; N13 C = +3.0 R 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 N18 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,
256.18 mbmp. (B) Undeformed dino£agellate cyst in cemented limestone from the upper slope. Spherical preservation of the cyst
implies early lithi¢cation. Clino, 497.89 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, 253.14 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 166-1007C-34R-3, 125^129 cm; depth 623.30 m. (E) Sample from dark layer shows foraminifer tests that collapsed
due to mechanical compaction. Interval 166-1007C-21R-5, 03^08 cm; depth 499.50 m. (F) Fine-grained, uncemented matrix of
dark, compacted layer. Note presence of coccoliths. Interval 166-1005C-23R-2, 19^23 cm; depth 591.39 m. (G) Detail of coccolith-rich matrix of dark layer with high micro porosity. Interval 166-1007C-21R-5, 3^8 cm; depth 499.50 m. (H) Light layer with
microsparitic cementation and moldic porosity. Interval 166-1003C-2R-1, 65^69 cm; depth 416.25 m.
39
40
2001b). Therefore, the change must re£ect
changes in the N18 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 N18 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.
4.4.2.1. 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.
4.4.2.2. 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.,
1997b; 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
41
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.
4.4.3. 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
42
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 lowstand meteoric 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
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).
5.2. 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 occur-
43
rence of marine isotopic values (N18 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.
5.3. 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
44
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 ( 6 4000 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 CO2 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 (443.49^452.94 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
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.
5.4. 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
45
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 lowMg 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
46
organic 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.
DSr ¼
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).
5.5. Strontium content as an indicator of diagenesis
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 in-
Sr=Ca ðmineralÞ
Sr=Ca ðfluidÞ
In the case of seawater the DSr for aragonite is
approximately unity, meaning that there is no active accumulation or discrimination of the element into this phase. In contrast the DSr for organic LMC is 0.12 and for inorganic LMC is 0.05
(Kinsman, 1969; Veizer, 1983). The DSr for dolomite is also about 0.016 (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 Ca2þ concentrations (50^60 mM) resulting in £uids with low Sr2þ /Ca2þ 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-
ues. Meteoric £uids percolating through a young
metastable carbonate terrain can also have highly
elevated Sr2þ /Ca2þ ratios, as the recrystallization
of the carbonates preferentially excludes Sr from
the calcite structure. Hence the notion of using
the Sr concentration of the diagenetic carbonate
as an indicator of the speci¢c type of environment
of diagenesis must be considered within other evidence.
An example of closed system diagenesis leading
to elevated Sr concentrations in diagenetic carbonates occurs within the pore £uids of periplatform
sediments deposited on the margin of Great Bahama Bank, where dolomites attained Sr concentrations as high as 2000 ppm (Swart and Melim,
2000) and celestite is common (Melim et al.,
2001b). Previous to such ¢ndings, concentrations
as high as these values would have been interpreted as re£ecting formation in an evaporite environment. Examples of the types of changes in
the Sr/Ca ratio of the pore £uids, which might be
expected as a result of closed system, are evident
in the data reported by Eberli et al. (1997a) and
Kramer et al. (2000). These studies of pore £uids
from sediments cored during Leg 166 show an
interval of relatively constant Sr concentration
to a depth of 50 mbsf overlying an interval in
which the Sr concentration steadily increases.
The absence of geochemical gradients indicates
either an absence of recrystallization or active circulation of seawater within this interval. Below
this upper £ushed zone, the concentrations of Sr
steadily increase as a result of carbonate recrystallization. The upper concentrations of Sr which
can be attained during this process is dictated
by the solubility of the mineral celestite (SrSO4 )
(Baker and Bloomer, 1988; Swart and Guzikowski, 1988). In the absence of signi¢cant sulfate
reduction a maximum of 600 WM of Sr could be
added to the pore £uids, which in turn would
produce calcite cements with Sr concentrations
of 3000 ppm. Higher concentrations of Sr in the
pore £uids, up to 4 mM, were reported in the
drilling of sediments on Leg 166 and are associated with the oxidation of organic material by
sulfate reduction.
Trends in the Sr concentrations of diagenetic
carbonates can be a clue to past hydrologic and
47
diagenetic regimes. A particularly good example is
the Sr measurements made on dolomites form
Clino below the hardground at 586.3 m. At this
locality Swart and Melim (2000) reported that the
concentrations of Sr increased from around 250
to 1300 ppm over a depth of about 100 m. However the ¢rst 20 m below the hardground surface
which represented a hiatus of about 2 Myr
showed an absence of signi¢cant geochemical gradients, in essence a fossil £ushed zone. The reduced thickness of this zone in comparison to
the modern £ushed zone suggests a di¡erent hydrological regime at this time. Swart and Melim
(2000) calculated the Sr/Ca in the pore £uids
which might have accounted for this increase
and determined that Sr concentrations in the lower dolomites predicted a reduction in the sulfate
concentration. A dramatic comparison to these
trends in the geochemistry of the dolomite below
the 586.3 m hardground in Clino is evident in the
100% dolomitized in Unda. In Unda the dolomites all have concentrations which might expected as a result of dolomitization from an
open system involving seawater.
Hence, as in the case of petrographic fabrics,
established truisms governing the concentration of
Sr in diagenetic carbonates have been challenged.
Relatively low concentrations of Sr in both dolomites and calcites can be produced in open-system
marine diagenetic environments, whereas high
concentration can be produced in closed-system
environments. Therefore, although freshwater environments can produce diagenetic cements with
low Sr concentrations, these values are not diagnostic.
6. Conclusions
The Bahamas transect o¡ered an unusual opportunity to study early diagenesis of aragoniterich carbonate sediments altered in a spectrum of
diagenetic environments from meteoric through
mixing-zone and into the marine-burial to deepburial environments. Although some overprinting
does occur, in general these di¡erent environments altered distinct portions of the Bahamas
transect, allowing each to be characterized. The
48
results question a number of current diagenetic
paradigms that have been broadly applied beyond
their original parameters. Speci¢c new ¢ndings
include:
(1) Large-scale sea level lowstands do not always have a chemically active meteoric lens. Examination of bank-top cores at the depth of the
latest Pleistocene lowstand revealed no evidence
of meteoric diagenesis. Rather the deepest meteoric diagenesis present appears to correlate to the
¢rst time each core location was subaerially exposed. In this case, the active meteoric lens extended W10 m below the land surface indicated
by a subaerial exposure horizon. Thus, any model
that uses meteoric diagenesis to track sea level
risks missing the large sea-level events.
(2) Meteoric-marine mixing zones in these cores
are characterized by aragonite dissolution and minor LMC cementation. Contrary to existing mixing-zone models, vuggy to cavernous porosity is
absent and dolomite does not occur. Current mixing-zone models are based on data from coastal
mixing zones where £ow rates are higher, surface
organic matter is common, and mixing relations
are complex. The inland position of Clino and
Unda during sea-level lowstands produces a different style of mixing-zone diagenesis ; one perhaps better suited to general application as the
coastal zone is very narrow compared to the complete extent of the mixing zone beneath a meteoric
lens.
(3) Marine-burial diagenesis mimics meteoric
diagenesis. Many diagenetic features thought to
be indicative of phreatic meteoric diagenesis also
occur in the marine-burial environment. These
features include blocky spar, moldic porosity,
neomorphism and microspar. The character of
marine-burial diagenesis is partially controlled
by the nature of the sediment being altered. We
have identi¢ed two end-member styles, an opensystem style characterized by dissolution of aragonite without signi¢cant cementation and a more
closed-system style with aragonite dissolution accompanied by calcite cementation. The similarity
of marine-burial and meteoric fabrics are caused
by selective aragonite dissolution in the shallow
marine burial realm shifting the pore water geochemistry towards the low Mg/Ca-ratios that are
also typical for freshwater. Such low ratios lead to
precipitation of typically shaped equant calcite
crystals (blocky spar, microspar). On the basis
of our ¢ndings, the current practice of routinely
assigning these fabrics to the meteoric diagenetic
environment has to be questioned. Without independent evidence of meteoric diagenesis (such as
caliche horizons, vadose fabrics or stable isotopic
data) these fabrics cannot be used to unequivocally identify any speci¢c diagenetic environment.
We further suggest that possibly some older studies based on the interpretation of these features
should be revisited.
(4) Microspar is a primary cement lithifying
aragonite-dominated mud. Microspar does not
form by aggrading neomorphism from recrystallized micrite. Rather, it forms as a cement between and around aragonite needles, that are later
dissolved. Microspar cementation can occur very
early in marine-burial diagenesis, prior to detectable mechanical compaction.
(5) Strontium content of sediments altered in
marine pore £uids can show an extreme range
of values, formerly thought to indicate di¡erent
environments. Very high values form in closedsystem marine-burial diagenesis as Sr released by
aragonite dissolution accumulates in the pore £uids. Relatively low values, on the other hand, occur during open-system marine-burial diagenesis
where the Sr is £ushed from the system. Sr values,
therefore, must by used with caution when interpreting diagenetic environment.
In conclusion, we would like to emphasize the
need for integrated studies that do not depend on
any single form of data (e.g. petrography, isotopes) as similar fabrics are produced by a number of di¡erent diagenetic environments.
Acknowledgements
We thank the shipboard party of ODP Leg 166
for their descriptions and assistance with sampling.
BDP cores Unda and Clino were collected with
funding from the United States National Science
Foundation under Grants OCE 891-7295 and
OCE 910-4294 to R.N. Ginsburg and P.K.S.,
from the Industrial Associates of the Comparative
Sedimentology Laboratory of the Rosenstiel
School of Marine and Atmospheric Sciences and
from the Swiss National Science Foundation.
Diagenesis research on the BDP cores was
supported by United States Department of Energy
grant DE-FG05-92ER14253 to G.P.E. and P.K.S.
and by the German Science Foundation (DFG)
grant We 2492 to H.W. and earlier Re 1051 to
John Reijmer. Additional work was supported by
JOI-USSAC grants to P.K.S. and G.P.E. We are
grateful to the many colleagues who contributed to
this study by collaborations and numerous discussions, including Robert N. Ginsburg (University
of Miami, Miami, FL, USA), G. Michael Grammer (Texaco EpP, Houston, TX, USA), Martin J.
Head (University of Cambridge, UK), and John
Reijmer (Geomar, Kiel, Germany). Reviews by A.
Immenhauser (Vrije Universiteit, The Netherlands), N.P. James (Queen’s Unversity, Canada)
and J.J.G. Reijmer (GEOMAR, Germany) are
gratefully acknowledged. The SEM micrographs
were acquired in the SEM Lab of Kiel University ^
many thanks to Christian Samtleben, Ute Schuldt,
and Werner Reimann for their support.
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Questioning carbonate diagenetic paradigms

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