Effects of early sea-floor processes on the taphonomy of temperate shelf skeletal carbonate deposits

Earth-Science Reviews 63 (2003) 1 – 31 www.elsevier.com/locate/earscirev

Effects of early sea-floor processes on the taphonomy of temperate shelf skeletal carbonate deposits Abigail M. Smith a,*, Campbell S. Nelson b b

a Department of Marine Science, University of Otago, P.O. Box 56, Dunedin, New Zealand Department of Earth Sciences, University of Waikato, Private Bag 3105, Hamilton, New Zealand

Received 1 May 2002; accepted 3 December 2002

Abstract Cool-water shelf carbonates differ from tropical carbonates in their sources, modes, and rates of deposition, geochemistry, and diagenesis. Inorganic precipitation, marine cementation, and sediment accumulation rates are absent or slow in cool waters, so that temperate carbonates remain longer at or near the sea bed. Early sea-floor processes, occurring between biogenic calcification and ultimate deposition, thus take on an important role, and there is the potential for considerable taphonomic loss of skeletal information into the fossilised record of cool-water carbonate deposits. The physical breakdown processes of dissociation, breakage, and abrasion are mediated mainly by hydraulic regime, and are always destructive. Impact damage reduces the size of grains, removes structure and therefore information, and ultimately may transform skeletal material into anonymous particles. Abrasion is highly selective amongst and within taxa, their skeletal form and structure strongly influencing resistance to mechanical breakdown. Dissolution and precipitation are the end-members of a two-way chemical equilibrium operating in sea water. In cool waters, inorganic precipitation is rare. There is conflicting opinion about the importance of diagenetic dissolution of carbonate skeletons on the temperate sea floor, but test maceration and early loss of aragonite in particular are reported. Dissolution may relate to undersaturated acidic pore waters generated locally by a combination of microbial metabolisation of organic matter, strong bioturbation, and oxidation of solid phase sulphides immediately beneath the sea floor in otherwise very slowly accumulating skeletal deposits. Laboratory experiments demonstrate that surface-to-volume ratio and skeletal mineralogy are both important in determining skeletal resistance to dissolution. Biological processes on the sea floor include encrustation and bioerosion. Encrustation, a constructive process, may be periodic or seasonal, and can be reversed. It produces both information and material. Bioerosion, in contrast, is destructive and permanent. In temperate areas bioerosion may destroy even very large shells during their long residence at the sea floor, on the order of hundreds to thousands of years. Overall, processes on the temperate sea floor may combine to destroy more carbonate than they produce, and the preservation potential of temperate shelf carbonate into the rock record may be significantly affected. Where preservation does occur in such a destructive regime, the effects of early sea-floor processes will be key determinants of the deposit, resulting in a ‘‘taphofacies’’ characteristic of temperate shelf carbonate sediments. D 2003 Elsevier Science B.V. All rights reserved. Keywords: Early diagenesis; Taphonomy; Cool-water carbonates; Abrasion; Breakage; Dissolution; Cementation; Encrustation; Bioerosion; Sediments

* Corresponding author. Tel.: +64-3-479-7470; fax: +64-3-479-8336. E-mail address: [email protected] (A.M. Smith). 0012-8252/03/$ - see front matter D 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0012-8252(02)00164-2

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1. Introduction Despite the increasing volume of literature on cooland cold-water carbonates (e.g., Nelson, 1988a; Rao, 1996; James and Clarke, 1997), geochemical and diagenetic scenarios for carbonate shelf deposits remain strongly biased towards warm-water, tropicalstyle perceptions. Even relatively recent texts discussing carbonate geochemistry and diagenesis reflect a specifically, though often unstated, tropical style (e.g., Morse and Mackenzie, 1990; Tucker and Bathurst, 1990; Tucker and Wright, 1990). Cool- or temperatewater carbonates, however, can differ markedly from their tropical counterparts in modes and rates of sediment formation, carbonate mineralogy and geochemistry, depositional rates, and therefore in their response to diagenetic processes (Table 1). Inorganic precipitation, extensive cementation, high production rates, and rapid burial of carbonate are all typical of early marine diagenesis (sensu Morse and Mackenzie, 1990) in the highly saturated warm waters of the tropics, and produce a constructive diagenetic sea-floor environment with high preservation potential for carbonate sediments (e.g., Bathurst, 1975). In con-

trast, the early marine diagenetic regime in temperate seas has commonly been labelled degradational or destructive (e.g., Alexandersson, 1979; Young and Nelson, 1988; Freiwald, 1998). Destructive diagenetic processes alter, dissolve, and remove grains, reduce grain size, damage or destroy structure, and generally eliminate information from the potential rock record. Early marine diagenesis in temperate environments seldom produces new grains or structures, and seldom adds to overall carbonate volume and information content. Merely having a mineralised skeleton is insufficient for preservation and/or fossilisation. Physical, chemical, and biological factors combine to determine the ultimate fate of skeletal material. Each destructive or constructive process interacts with the results of other factors, as when, for example, endolithic borings weaken mollusc shells which are then more susceptible to physical destruction by abrasion (Young and Nelson, 1988) or chemical loss by dissolution (Canfield and Raiswell, 1991). When these destructive forces interact to remove material they often do so in a selective way, thus altering the composition of the assemblage available for preservation. Such tapho-

Table 1 Environmental, compositional, and diagenetic features for typical tropical and temperate shelf carbonate facies (adapted from Nelson, 1988b; James, 1997) Process

Tropical—subtropical

Temperate—polar

Self morphology

Rimmed shelf (high to low energy)

Unrimmed shelf or ramp (high energy)

Carbonate saturation level

Supersaturation

Supersaturation to undersaturation

Carbonate factory

In photic zone (inner shelf depths)

Mainly in aphotic zone ((inner-)mid-outer shelf depths)

Skeletal associations

Chlorozoan (corals, calcareous green algae, molluscs, benthic forams)

Foramol (bryozoans, echinoderms, bivalve molluscs, forams)

Nonskeletal carbonate grains

Ooids, aggregates and/or peloids common

Absent

Sedimentation rate

High carbonate accumulation rate (>10 cm/kyr)

Low carbonate accumulation rate ( < 10 cm/kyr)

Carbonate mud production (source)

Common to abundant (floral disaggregation, inorganic precipitation)

Rare, may be locally important (mechanical abrasion, bioerosion, microfossils)

Primary mineralogy

Aragonite plus high-Mg calcite

Low-Mg calcite plus intermediate-Mg calcite

Sea-floor cementation

Locally common, sometimes pervasive

Uncommon

Marine diagenetic regime

Mainly constructive (grain preservation and chemical precipitation)

Often destructive (grain abrasion, biodegradation, and dissolution)

Shell preservation

Thick shells, good preservation potential

Sometimes thinner shells, poor to good preservation potential

Overall preservation potential

High

Low

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nomic processes occur mainly at the sediment – water interface and in the bioturbated layer just below (Davies et al., 1989a). Where sedimentation rates are high, as in most tropical carbonates, skeletal material will be buried rapidly and thus escape prolonged exposure to sea-floor processes. Chave (1964) performed a set of experiments using tropical carbonates from Bermuda and Campeche Bank in which he summarised the paleoecological taphonomy of tropical carbonates. Here we synthesise a number of recent laboratory and field investigations into a review of the effects of early sea-floor processes on cool- and cold-water carbonate sediments. These processes may affect skeletal carbonate any time after calcification, but generally their effect is greatest after death. Physical processes of abrasion and breakage associated with transport, chemical changes due to interactions with ambient sea water, and biologically mediated processes may occur while sediment is at or near the sea floor. Because carbonate precipitation, cementation, and sedimentation are all rare or slow in

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cool waters, temperate carbonate sediments such as those mantling extensive areas of the southern Australian and New Zealand shelves may remain in this seafloor zone for hundreds to many thousands of years (e.g., Nelson et al., 1988; James, 1997). Consequently, early sea-floor processes in temperate carbonates can become of considerable importance, affecting both preservation potential and the eventual character of preserved material (Fig. 1).

2. Physical breakdown 2.1. Description of physical breakdown processes Any grain that is transported in water, even for a very short distance, is subject to the mechanical forces of gravity and friction. Most physical movement, especially in the shallow and shelf marine environment, is due to water energy, though there may be some movement caused by intermittent forces such as macro-

Fig. 1. Taphonomic ‘filters’ active in limiting preservation of shelf carbonate sediments, with particular reference to early sea floor.

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organism disturbance or seismic activity. Dissociation and breakage are the first stages of mechanical breakdown, especially for large skeletons. Articulated organisms such as barnacles, bivalves, echinoderms, and brachiopods fall easily into their component parts on decay of the organic material that held them together (Allison, 1986; Kidwell and Baumiller, 1990; Daley, 1993), thus providing smaller units for subsequent processes to affect. Other types of skeletons (e.g., certain molluscs) may fracture along zones of weakness into pieces or unit crystals (Force, 1969), strongly related to the shape and strength of the skeleton. A delicately branched bryozoan colony, for example, loses its branches after death, forming Y-shaped fragments, whereas a foliose colony breaks into roughly rectilinear plates (Bone and James, 1993; Smith and Nelson, 1996). The Sorby Principle—defined by Folk and Robles (1964) after Sorby (1879) that some inherent structural property of skeletons influences fragment size—can be extended to include fragment shape and perhaps breakdown rate. Broken pieces or small skeletons begin to act as single sediment grains and become susceptible to attrition and abrasion. Movement in water alone is not particularly abrasive, but bouncing off other solid particles, particularly hard siliciclastic grains, can cause considerable damage (Chave, 1964). In the transport processes of saltation, rolling, and entrainment, skeletal grains become increasingly abraded. Terrigenous material associated with skeletal carbonate sediment, a common situation on many temperate shelves (Nelson, 1988a), dramatically increases the abrasion rate of carbonate grains (Shroba, 1993). An abraded skeletal grain has rounded edges and a characteristic smooth shiny surface. Ornamentations such as spines and ridges are missing and pores may be enlarged. Some of the effects of abrasion on bryozoan particles, a dominant skeletal component of many temperate shelf carbonates (Nelson, 1988b; Bone and James, 1993; Hayton et al., 1995), are illustrated in Fig. 2. Erosion and blurring of ornamentation and surface lamellae is initially rapid, and then becomes more gradual. On platy or elongate grains, rounding and polishing is most extensive on edges; the centre of the grain may be better preserved (Perry, 2000). The nature of highly abraded skeletal fragments is frequently difficult to determine.

2.2. Intrinsic controls on physical breakdown Chave’s (1964) now-classic tumbling experiments found major differences in resistance to breakage and abrasion among marine taxa. Compact snail shells were most durable; among the least durable were calcifying algae and branching bryozoans (Table 2). Robust corals and bryozoans, and weaker molluscs were intermediate in resistance. Tumbling experiments comparing various bryozoan colonial growth forms have shown that, even within a phylum, different morphologies respond quite differently to the same abrasion pressure (see Fig. 5; Smith and Nelson, 1996). Branched and encrusting bryozoans were destroyed by physical abrasion fairly rapidly, the most susceptible being the multilaminar species Celleporaria agglutinans, which lost about 50% of its original weight after 127 h tumbling. Compact free-living Otionella spp. and erect fenestrate forms, on the other hand, were more resistant, retaining more than 80% of their original weight after 127 h tumbling. Different size classes of the same species may well act quite differently; small specimens of the bivalve Spisula sp. were far more breakable than large ones (Chave, 1964). Even within a single skeleton, response of valves to abrasion stress may differ, for example, in brachiopods the pedicle valve is more robust than the brachial (Daley, 1993). If marine taxa do break down into characteristic shapes and sizes (following the Sorby Principle), this taxonomic effect is really secondary (e.g., Hoskin et al., 1983). Composition, size, and shape are the primary intrinsic controls on resistance to abrasion. Carbonate particles are destroyed by abrasion much faster than siliciclastic grains—in one experiment limestone pebbles lasted less than a quarter as long as quartzite pebbles (Sneed and Folk, 1958; see also Matthews, 1983). Among calcareous bioclasts, however, carbonate mineralogy seems to have little effect on resistance to abrasion (Chave, 1964; Smith and Nelson, 1996). Microarchitecture and content of organic matrix are more important (Table 2). Most resistant carbonate grains are finely crystalline and compact with relatively little organic matrix (strong molluscs). Brittle or porous organic-rich skeletons with coarsely crystalline texture have moderate resistance to abrasion (corals, fragile molluscs). The most fragile skeletons are open-

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Fig. 2. Effects of early sea-floor processes on temperate shelf skeletons. Fresh temperate skeletal carbonate was subjected to abrasion (Smith and Nelson, 1996) and dissolution (Smith et al., 1992) in laboratory settings, and to biological processes in situ (Smith, 1993), for comparison with an ‘old-looking’ sediment grain.

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Table 2 Resistance of various skeletal components to abrasion and breakage (data from Chave, 1964; Smith and Nelson, 1996) Durable

Intermediate

Fragile

Taxa

Large bivalves (oysters, clams) Compact snails (Nerita)

Small bivalves (scallops, Anomia) Weak gastropods (limpets, abalone) Robust corals Free-living bryozoans

Echinoids and starfish (Poaster, Stronglocentrotus) Coralline and codiacean algae (Corallina, Penicillus) Delicate corals Branching and encrusting bryozoans

Texture

Dense, compact Fine-grained

Hard, brittle Coarse-grained

Open, branching Coarse-grained

Composition

Little organic matrix

Average organic matrix

Much organic matrix

structured and organic-rich (calcareous algae, branching bryozoans). Grain size is a major control on susceptibility to abrasion. Large grains lose volume more rapidly than small ones because of their greater surface area. Large grains may be closer to the original skeleton in shape, and may therefore have more ornamentation available for removal. In general, the rate of abrasion is greater in grains with a high surface-area-to-weight ratio, especially in a poorly sorted deposit (Martin, 1999). Heavier shells, however, tend to be buried more rapidly than lighter ones, thus escaping the abrasive regime of the sea floor (Driscoll, 1970). Shape, as well as composition and size, affects the longevity of bioclastic particles; different shapes have different hydraulic properties. Large plate-shaped grains, for example, settle four times more slowly than their theoretical settling velocities (calculated from density), whereas large block-shaped grains settle nearly at calculated rates (Maiklem, 1968). Because settling velocity is related to shape as well as density, it can be assumed that critical entrainment velocity would be similarly variable. Thus, Nelson and Hancock (1984) demonstrated hydraulic equivalence amongst differently sized skeletal taxa in flume experiments: a current of 20 cm/s can simultaneously transport medium sand-sized quartz grains, coarse sand-sized platy bivalve fragments, very coarse sand-sized bryozoan grains, and granule-sized echinoderm plates. Other studies have shown that stage of organic decay (Allison, 1986; Daley, 1993), degree of previous shell damage (Young and Nelson, 1988), and encrustation or cementation (Mitchell-Tapping, 1981) can also affect shell abrasion rates. It is the sum of all these intrinsic properties that determine the response of a skeletal grain to an abrasive regime.

2.3. Extrinsic controls on physical breakdown The primary external control on breakage and abrasion is hydraulic regime. In very shallow water, waves are most active, reaching transport rates on the order of 1 m/s (Collins, 1988; Kukal, 1990), but in deeper shelf waters bottom currents (typically 0.1– 1.0 m/s) are the major agents of transport (e.g., Bone and James, 1993). Abrasion is greatest in the surf zone (Driscoll, 1967; Force, 1969), though episodic climatic events can also have dramatic effects (Davies et al., 1989b). Mode of transport is also relevant, as bedload transport is more abrasive than travelling in suspension (Martin, 1999). A second effect is the degree of terrigenous sediment mixing. The more siliciclastic ‘‘grit’’ within a carbonate deposit, the greater the potential attrition rate of associated skeletal grains (Shroba, 1993). A feature of many temperate shelf carbonate facies is their mixed siliciclastic – skeletal nature (Nelson et al., 1982; Nelson, 1988a; Gillespie and Nelson, 1997). In general, abrasion rates rise with increasing grain size (Powell et al., 1989), itself a function of strengthening hydraulic regime, although sand and small pebbles have been shown to be more abrasive than large pebbles (Driscoll, 1967; Force, 1969), while mud and clay are generally not very abrasive (Aguirre and Farinati, 1999). Poor sorting enhances abrasion (Driscoll and Weltin, 1973). The third extrinsic control on physical breakdown is exposure time. In some models time is replaced by distance, which takes into account the episodic nature of transport processes. In some regimes, however, a grain may not lie far from its point of origin but may have rolled back and forth over long distances. Such oscillation of grains is one reason that beach surf is the most abrasive of environments (Kukal, 1990); deeper water currents are both slower

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and less oscillatory. A single grain may move up to 3000 m/day in the surf zone but perhaps only 150 m/ year in shallow tidal sand waves (Menard and Boucot, 1951). Aguirre and Farinati (1999) found that time of exposure to surf zone dynamics is a better predictor of ‘‘abrasion signature’’ than either distance or total time in transit. The relatively slow sedimentation rates typical of cool-water carbonate shelves, on the order of only a few cm/kyr (Fig. 3), mean that shell material may remain exposed to abrasive forces for very much longer than in many tropical shallow-water settings where rapid carbonate production and sedimentation, as well as possibly early cementation, tend to protect skeletons from abrasion (Pilkey et al., 1979).

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2.4. Conceptual model and rates of physical breakdown Physical alteration of carbonate grains by dissociation, abrasion, and breakage is a cumulative, oneway, destructive early sea-floor process (Fig. 4). The effects of reducing grain size, removing structure, and reducing information available are irreversible and ongoing. As long as a carbonate grain is at or near the sediment – seawater interface within the taphonomically active zone of Powell et al. (1989), it is susceptible to the physical destruction brought about by transport. If the physical regime is severe enough, or lasts long enough, grains may be ground down over time to an anonymous carbonate mud, which may be

Fig. 3. Estimates of Quaternary rates of production and accumulation (cm/kyr) for many warm- and cool-water shelf carbonates, along with sedimentation rates for Tertiary Australasian temperate limestones and many other global carbonate deposits (adapted from Smith, 1988 and James, 1997).

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Fig. 4. Conceptual model of the effect of mechanical sea-floor processes of abrasion and breakage on skeletal carbonate. Over time, skeletal material decreases in size and morphological details are obscured, reducing environmental information available.

bypassed to deep water or destroyed by chemical diagenesis. Since skeletal structure strongly influences breakdown, the biotic composition of carbonate producers is a critical determinant of susceptibility of carbonate sediment to abrasion pressure (e.g., Martin, 1999). Abrasion of an individual grain can be characterised by Sternberg’s Law, a simple exponential function: Ws = W0e as, where W0 is the original weight of the particle, Ws is the weight after transport over distance s, and a is a constant rate of abrasion (Kukal, 1990). Distance (s) is not the net distance travelled, but the total distance travelled over the life of the grain, and so could be replaced by time (t). The constant a depends on two factors: E for the extrinsic abrasive pressure of the hydraulic regime (including the degree of mixing with other sediments), and I for the intrinsic resistance of the skeletal grain. E is thus primarily a function of water velocity and proportion of siliciclastic material around the grain. I is a function of size, shape, composition, microstructure, and organic matrix of the grain. Thus, quartzite particles in a low-energy regime would have small I and E, whereas carbonate particles in the surf zone would have large I and E. It is the interplay of the factors extrinsic and intrinsic

to the skeleton that determines the extent of abrasion of a particle, and thus its ultimate preservation potential, as indicated by the model: Wt = W0e at where a = f(I,E) (Fig. 4). Laboratory experiments on abrasion are usually carried out in rock tumblers (e.g., Chave, 1964; Driscoll, 1970; Argast et al., 1987; Broadhead and Driese, 1994). Wave tanks and shakers have also been tried (e.g., Driscoll and Weltin, 1973; Mitchell-Tapping, 1981). Surf barrel experiments have shown 2– 17% loss in 500 h (Mitchell-Tapping, 1981) and 10– 80% loss in 1800 h on temperate beaches (Driscoll and Weltin, 1973). While these artificial abrasion regimes may reveal relative resistance, they are quite unlike natural settings (Kukal, 1990) and cannot be translated into real rates (Powell et al., 1989). There are, however, few data available on rates of carbonate abrasion in natural sedimentary environments. Very few of the laboratory simulations have gone so far as to calculate the rate of loss of carbonate material. A few studies, however, provide enough information to allow rough calculation. Using data from Driscoll (1967), for example, if 2 h in a rock tumbler is equivalent to 100 h in the surf zone, and bivalves lose 3 – 4% of their original weight (0.1 –0.5 g) in gravelly sand, assuming population density is on the

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order of 10 bivalves/m2, then carbonate lost through abrasion comes to 80– 400 g CaCO3/m2/year. 2.5. Sedimentary implications of physical breakdown An abraded sediment is missing information; the destructive forces of impact remove material and therefore fine structure from carbonate grains. Ultimately abrasion transforms recognisable skeletal gravels to much finer, featureless carbonate particles (Fig. 2). Abrasion produces finer-grained and better-sorted sediments, the side effects of hydraulic regime. Increasing the surface-to-volume ratio of skeletal carbonate increases the potential effects of chemical diagenesis and degradation, including maceration in undersaturated waters (Alexandersson, 1979) and enhanced susceptibility to dissolution and pressuredissolution phenomena. Abrasion may be an important source of carbonate mud, formed from small flakes of carbonate chipped off larger grains. Carbonate muds are reported from a variety of cool- and cold-water localities, including north Atlantic (Fitzgerald et al., 1979; Farrow and Fyfe, 1988), southeast Australia (Blom and Alsop, 1988), and the Black Sea (Trimonis, 1974). Such muds cannot be formed from disarticulation of codiacean green algae, the principal source of fine carbonate in tropical settings, and it has been suggested that

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bioerosion, shell abrasion, and planktonic tests are the principal contributors to temperate carbonate muds (Table 1; Farrow and Fyfe, 1988). Smith and Nelson (1996) found in short-term tumbling experiments that abrasion of different bryozoan growth forms quickly produced small to substantial quantities of carbonate mud, reaching as much as 40 –50% of starting weight after 127 h of tumbling in the case of multilaminar encrusting types (Fig. 5). Shell abrasion has been shown to be highly selective; microstructure, skeletal form, and therefore taxon all affect abrasion resistance (Chave, 1964; Smith and Nelson, 1996). Selectivity in a destructive process leads to taphonomic bias, because preservation potential is non-random. Non-representative preservation of a high-energy biotic assemblage should be anticipated, in which fragile taxa such as calcareous algae and bryozoans are missing (e.g., see Fig. 11). While it may be a problem in (paleo)ecologic reconstruction, evidence of abrasion nevertheless can enhance interpretation of hydraulic regimes (e.g., Pilkey et al., 1979). Physical and morphological evidence of high environmental energy levels would be strongly supported by the absence of forms susceptible to abrasion. In warm waters, a number of factors tend to protect skeletal sediments from extensive abrasion and breakage. For example, reef buildups and expansive tracts

Fig. 5. Abrasion resistance and carbonate mud production for different bryozoan growth forms in tumbling experiments (from Smith and Nelson, 1996). Key to bryozoan symbols in Fig. 11.

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of sea-grass beds may protect lagoonal sediments from transport and thus from appreciable abrasion, much as interstitial carbonates are sheltered within reef structures (Bathurst, 1975). Rapid burial and sedimentation (Fig. 3; Parsons-Hubbard et al., 1999) along with sea-floor cementation remove grains from the transport zone and minimise abrasion effects. Consequently, some rather fragile skeletons may be preserved in these tropical settings. Temperate carbonates are not usually so well protected. They are typically associated with high energy open shelves, mainly lacking reefs, although oyster mounds (Nelson et al., 1988) and coralline algal buildups (Freiwald and Henrich, 1994) can occur locally and, while kelp forests can be important in shallow waters off rocky coasts (e.g., Nelson et al., 1988; Freiwald, 1993), extensive tracts of sediment-baffling sea grasses are rare. Slow accumulation rates (Fig. 3) and minimal sea-floor cementation mean grains lie at or near the sediment – seawater interface for a long time and may experience regular, or periodic, abrasion. The locus of carbonate formation is typically in much deeper shelf zones in cool-water compared to tropical carbonates (James, 1997), and thus skeletal grains are not so susceptible to surf-zone dynamics. Cool-water shelves are very likely to include some siliciclastic material (Nelson, 1988a), thus increasing the force of abrasion over an all-carbonate setting. On the other hand, cooltemperate decay of organic matrices is slower than in the tropics, which may slow initial disarticulation and breakage (Kidwell and Baumiller, 1990). Some of the dominant taxa on temperate shelves (e.g., large bivalve molluscs) are more robust than the corals and codiacean algae dominating tropical carbonates. Many cool-water carbonate deposits show clear evidence of abrasion and winnowing. A typical example occurs on the Australian Otway Shelf (Boreen et al., 1993). Indeed, eustatic sea-level changes may allow abrasive hydraulic regimes to rework the entire shelf sediment blanket over time (Boreen et al., 1993). The rounded, polished, and often stained skeletal grains so common on temperate mid-shelves are probably mainly relict from reworking during late Quaternary oscillations of sea level (e.g., Nelson et al., 1982; Boreen and James, 1993; James, 1997). Chave (1964) described a typical high-energy deposit

as a coarse sandstone with only a few fossils: robust gastropods and fragments of less-robust bivalves. As a cool-water example, cross-bedded units of coarsegrained Abrakurrie Limestone, South Australia, contain only the most robust, rooted forms of bryozoans along with echinoid fragments and benthic foraminifera. They have been interpreted as high-energy open-shelf deposits subject to prolonged abrasion (James and Bone, 1991).

3. Dissolution and precipitation 3.1. Description of dissolution and precipitation The chemical equilibrium of calcium carbonate in sea water is a balance between dissolution and precipitation. Dissolution turns solid CaCO3 into its component ions, CO32 and Ca2 +, removing skeletal material, and therefore decreasing both grain size and the biotic information available for preservation. Inorganic precipitation creates new solid CaCO3 from ions in sea water but provides no information about organisms or their living environment; it may, however, provide important clues as to diagenetic environment. Such cementation also binds existing skeletal grains together, limiting their susceptibility to destruction by other processes. Dissolution of calcium carbonate in temperate shelf bottom waters has long been considered unusual and unexpected, because shallow waters are generally oversaturated with respect to calcite and aragonite worldwide (Rao, 1996). Observations of dissolution effects in shallow-water sediments have been puzzling and variously explained (e.g., Pytkowicz, 1965; Alexandersson, 1978; Lewy 1981; Aller, 1982; Nelson and Bornhold, 1983; Walter and Burton, 1990; Freiwald, 1998). It is only recently that the role of organic matter in sediments has come to the fore (e.g., Wollast, 1994). Even if overlying water is supersaturated, the microbial respiration of organic matter in sediments increases CO2 levels, acidifying porewater, and titrating CO32 sufficiently to (begin to) dissolve at least metastable carbonate phases such as Mg calcite and aragonite. This effect is greatest in aerobic sediments, and exacerbated by the effects of bioturbation (Aller, 1982). The oxidation of sulfides under the influence of high organic activity, too, produces

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undersaturated marine pore/bottom water (Morse and Mackenzie, 1990; Wollast, 1994). In oxygenated sediments, undersaturation with respect to calcium carbonate may occur only a few cm below the surface (Wollast, 1994). Dissolution at the sediment – seawater interface may not be the norm in temperate waters, but it clearly occurs (Cutler and Flessa, 1995), especially given the long sea-floor residence time experienced by many cool-water skeletal carbonate sediments (Fig. 3). Evidence for (near) sea-floor dissolution of carbonate skeletons has been described from many temperate oceanic shelf waters, including the NW Atlantic (Aller, 1982), NE Atlantic (Freiwald, 1998), SE Pacific (Pytkowicz, 1965), NE Pacific (Smith, 1971; Nelson and Bornhold, 1983), and SW Pacific (Nelson et al., 1988). The studies of highly degraded and fragile skeletons forming carbonate deposits in seasonally carbonate-undersaturated waters of the Skagerrak and Baltic Seas by Alexandersson (1972, 1975, 1978, 1979) and Lewy (1975, 1981) are particularly informative. Alexandersson (1975) recorded a progression of carbonate grain alteration, which he termed maceration, from surface grain etching about crystal heterogeneities to deeper penetration carbonate dissolution resulting in loss of ornamentation, the separation of subskeletal units, and eventually skeletal grain disappearance. He invoked carbonate undersaturation as the driving mechanism for carbonate degradation, while Lewy (1981) additionally emphasised the important role of early decomposition by microbes in the maceration process. Initially then, dissolution results in a ‘‘cleaning effect’’ of grains, where any internal cements and adhering or infilling carbonate detritus is removed. Rounding, surface etching, corrosion of margins, and loss of fine structure leads to a characteristically chalky, etched, and holed specimen (Alexandersson, 1975; Flessa and Brown, 1983) (Fig. 2). Ultimately, biogenic carbonate grains disintegrate into their component crystallites (Alexandersson, 1978; Henrich and Wefer, 1986). Tiny crystallites are prone to dissolution, especially in cold sea water, where the component ions become part of the solution load and thus available for other processes. Given that dissolution can occur in temperate shelf carbonates, does this mean that carbonate precipitation does not? Inorganic precipitation of calcium carbonate

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grains and cements is widespread and common in warm tropical sea water (Bathurst, 1975). The physico-chemical conditions on temperate shelves of cooler water temperatures, a lesser degree of saturation, and higher CO2 content, however, are much less likely to promote precipitation. Ooids, aggregates or grapestones, and other inorganic particles do not occur in temperate carbonates (Table 1). Neither is sea-floor cementation common in temperate carbonate sediments. Intraparticle high-Mg calcite spar fringes are reported in the chambers and pores of some skeletons from Mediterranean deposits (Alexandersson, 1974), attributed to biochemical activity of calcareous red algae. Rao (1981) also recorded sparry crystal precipitates (low-Mg calcite) inside certain bryozoan zooecia in Tasmanian shelf carbonates, but it is unclear whether the affected grains are truly modern or, as is more likely, reworked and relict. Some high-Mg calcite cementation of grains has been observed in the Kattegat Sea (latitude 58jN), related to sea water mixing with bicarbonateenriched meteoric water (Jørgensen, 1976). But most temperate shelf carbonate sediments lack any interparticle cements and remain loose and free-moving (e.g., Young and Nelson, 1988). Nevertheless, marine cementation by intermediate-Mg fibrous spar and micrite has produced hardground-like slabs of lithified limestone locally on some temperate shelves, such as near the shelf-slope break off southern Australia (e.g., Nelson and James, 2000). These authors suggested that high energy conditions and prolonged seawater pumping through the bottom sediments during times of lowered sea level and very slow or no sedimentation for long periods of time were prerequisites for sea-floor cementation of any significance to occur in temperate carbonates. Development of near-sea bed carbonate nodules and concretions (Mg calcitic, protodolomitic) also occurs in temperate waters, often in more siliciclastic-dominated settings, where it appears to be mainly associated with shallow burial zones of microbial degradation of organic matter (Irwin et al., 1977), instigated by sulphate-reducing bacteria or oxidation of methane (e.g., Nelson and Lawrence, 1984; Farrow and Fyfe, 1988). Even when taken together, the above examples of carbonate precipitation at the shallow temperate sea floor remain exceptions rather than the rule. The calcitic nature of most cool-water carbonate sedi-

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ments, and their correspondingly low diagenetic potential, mean that their lithification into limestones typically is delayed until deep burial when pressuredissolution processes can develop (Nelson et al., 1988; Nicolaides and Wallace, 1997; Nelson and James, 2000). 3.2. Controls on dissolution Not all carbonate particles have the same resistance to dissolution. While composition, particularly carbonate mineralogy, plays a part, structural parameters that influence the surface area available for dissolution are probably more important (Walter and Morse, 1985). Reactive surface area available for dissolution varies with both grain diameter and grain shape, and, to a more limited extent, porosity and surface ornamentation (Walter and Morse, 1984). Reduction in grain size leads to enhanced potential for dissolution due to increased surface-to-volume ratio. Gross morphology of skeletal particles is critical in determining susceptibility to dissolution (Flessa and Brown, 1983; Smith et al., 1992): robust and compact shapes are more resistant than delicate porous branches. Microstructure is also important (Walter and Morse, 1984), as surface roughness changes reactive surface area. Persistence of organic coatings, shapes and sizes of component crystallites, and porosity and permeability of a skeleton all affect reactive surface area and thus dissolution rate. For example, the red alga Amphiroa fragilissima with a

lamellar structure of Mg-calcite needles of high porosity and permeability, dissolves far more rapidly than Marginopora vertebralis, a benthic foraminifer with similar Mg-calcite arranged in rods of low porosity and permeability (Henrich and Wefer, 1986) (Table 3). Mineralogy has been shown to be a poor predictor of dissolution, at least amongst bryozoans (Smith et al., 1992). Surface waters are less saturated with respect to aragonite than with calcite, and aragonite has a lower thermodynamic stability. Thus, aragonite should be more likely to dissolve in any given conditions than is calcite (Chave et al., 1962). In fact, aragonitic shells do sometimes turn chalky in modern temperate settings (e.g., Nelson et al., 1982, 1988; Boreen and James, 1993; Gillespie and Nelson, 1997) and appear to have been preferentially removed before lithification of many temperate limestones (e.g., Beu et al., 1972; Nelson, 1978). Mgcalcite, too, may leach Mg over time in cool sea water and trend towards stable low-Mg calcite (Nelson et al., 1988). Increasing mol% of Mg in calcium carbonate increases its solubility, particularly over the range of 4 –16 mol% MgCO3 (Wollast, 1994). While Mg calcite may dissolve faster than pure calcite, the cumulative effects of structure and morphology are more important than such small mineralogical differences. Experimental dissolution studies have shown that, in biogenic carbonates, reactive surface area is the most important factor. Flessa and Brown (1983) found the dissolution rate of macroinvertebrate carbonate

Table 3 Resistance of various skeletal components to dissolution (data from Flessa and Brown, 1983; Walter and Morse, 1985; Henrich and Wefer, 1986; Smith et al., 1992) Resistant

Intermediate

Susceptible

Large bivalves (cockles, oysters) Compact snails, opercula (Crucibulum, Turbo) Echinoids (Encope)

Small bivalves (Tagelus, Lucina)

Free-living, massive bryozoans

Red algae (Amphiroa) Corals Fenestrate, encrusting bryozoans

Delicate bivalves (Felaniella) Barnacles (Cthalamus) Benthic foraminifera (Marginopora) Codiacean green algae (Halimeda) Delicate branching bryozoans

Texture

Little ornamentation Low porosity Large, compact

Some ornamentation Moderate porosity Moderate

Extensive ornamentation High porosity Small, delicate

Composition

Calcite

High-Mg calcite

Aragonite

Taxa

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skeletons to be strongly related to their weight and surface area (and therefore density), rather than simply mineralogy. Among 16 species studied in their dissolution experiments, dense low-Mg calcite oyster shells were the slowest to dissolve (losing 2 –3%/h), while porous barnacle plates having the same low-Mg calcite mineralogy were 20 times more soluble than the oysters (losing 90%/h). Walter and Morse (1985) re-emphasised that the microstructure of skeletal grains can be a more significant control on shell dissolution rates than thermodynamic (mineralogic) stability. Laboratory simulation of bryozoan colony dissolution (Smith et al., 1992) has shown that morphology interacts in a complex way with mineralogy to determine preservation potential amongst bryozoans (Fig. 6). At any given time, the balance between dissolution and precipitation will be determined by a variety of physico-chemical factors associated with sea water, particularly CaCO3 saturation level (which is in turn linked with temperature, pressure, salinity, and pH)

Fig. 6. Rates of dissolution of different bryozoan growth forms in relation to mineralogy in acid bath experiments (from Smith et al., 1992). Key to bryozoan symbols in Fig. 11.

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(see e.g., Chave and Schmalz, 1966; Burton, 1993). Since the continental shelf is well above the lyscocline, water depth has little effect on dissolution. Numerous studies have shown a clear relationship between dissolution rate and seawater saturation properties (e.g., Keir, 1980; Walter and Morse, 1984, 1985). There are less obvious effects on dissolution exerted by living organisms. Some bivalves dissolve their own shells using biogenic acids when they become anaerobic. They use carbonate dissolved from inside the shell (mainly inside the pallial line) as an alkali reservoir to neutralise acid produced during oxygen deprivation. The process is presumably reversed when environmental stress is released (Crenshaw, 1980). Bioturbation, too, can increase dissolution in sediments (Aller, 1982; Walter and Burton, 1990). 3.3. Controls on precipitation Sea-floor cementation occurs when there is a sufficiently large source of CaCO3 in solution (i.e., a high degree of supersaturation) and an appropriate environment for precipitation (Nelson and James, 2000). In temperate shelf settings, time is probably important (Fig. 3). Supersaturation alone is insufficient. The presence of organic material, particularly humic compounds, can inhibit precipitation even in highly supersaturated sea water (Berner et al., 1978). Cementation is not strongly related to grain-size or reactive surface area of sediments. The mineralogy of cement reflects an interplay of kinetic and thermodynamic factors. Cements may also be influenced by substrate, hydrology, and organic compounds present. The rare sea-floor cements that occur in cool- and cold-water carbonates are usually low- to intermediate-Mg calcites (up to 12 mol% MgCO3), rather than the more common aragonite or high-Mg calcite (>12 mol% MgCO3) cements found in tropical submarine cementation (e.g., Rao, 1981; James and Bone, 1992; Nelson and James, 2000). These non-ferroan Mg-calcitic cements are usually precipitated in skeletal pores and interstices, and less commonly may form external rims about grains, binding them together. Precipitation is commonly biologically mediated, such as the infilling of conceptacles in coralline algae with Mg-calcitic spar (Moberly, 1970).

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3.4. Conceptual model and rates of dissolution and precipitation The chemical equilibrium of calcium carbonate in sea water occupies a spectrum of possibilities with two end members: one mainly constructive, one mainly destructive (Fig. 7). The response of skeletal grains to physico-chemical seawater factors is related to mineralogy and morphology. Dissolution is not truly reversible—only the volume of carbonate can be replaced by precipitation; the environmental information in the original skeletal carbonate is lost. Dissolution rate R (in Amol/g/h) is a function of saturation state of seawater (X), a dissolution rate constant (k), which is related to reactive surface area of the grain, and the order of the reaction (n), which is related to mineralogy. The equation R = k(1 X)n indicates that dissolution varies linearly with saturation state and mineralogy, but at different slopes depending on the reactive surface area of grains (Walter and Morse, 1984). Saturation state (X) is the ratio of the

total ion concentration product ([Ca2 +][CO32 ]) to the apparent solubility product of the mineral in question (Ksp V ) (Keir, 1980). Karagonite V = 6.65  10 7, 7 while Kcalcite V = 4.39  10 (Morse et al., 1980). In experiments, X is commonly kept between 0.26 and 3.98 (e.g., Walter and Morse, 1985). The rate constant (k) varies considerably, from about 50 to almost 2000 Amol/g/h (Walter and Morse, 1984). This high degree of variability reflects variation in surface roughness and size of grains. Experimentally determined surface roughness (sr) varies from rhombic calcite (sr = 1) to Halimeda fragments (sr = 6.9). Reaction order is a constant, independent of species or taxon. Walter and Morse (1985) found naragonite to be about 2.5, whereas Keir (1980) calculated it at 4.2. Similarly, Keir’s (1980) ncalcite is 4.5, whereas Walter and Morse (1985) found it to be 2.9. Walter and Morse’s calculations, based on shallow-water invertebrates, are more usefully applied to the temperate shelf, as Keir’s were on planktic microfossils in the deep ocean. Walter and Morse (1985) thus calculated theo-

Fig. 7. Conceptual model of the effect of chemical sea-floor processes of dissolution and precipitation on skeletal carbonate. Dissolution effects dominate overwhelmingly, but selectively, in cool-water shelf carbonates, so that over time certain skeletal materials lose size, coherence, and identity, even to the point of being totally dissolved.

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retical dissolution rates on the order of 50 – 150 Amol/ g/h, depending on the pH of sea water. Alternatively, a simpler model might be an adaptation of Sternberg’s law. The weight (W) of a grain after time (t) is the original weight (W0) less dissolution (a function of grain characteristics (a), water saturation state (X), and time) plus cementation (a function of water saturation state and time): Wt = W0 f(a,X,t) + f(X,t). Real-life dissolution rates can vary dramatically. In warm tropical waters, back-reef carbonate particles may dissolve at rates varying from 0.001 g/year (corals, echinoids) to 0.05 g/year (red algae) (Walter and Burton, 1990). In contrast, mussels on the Louisiana slope may dissolve at up to 0.1 g/year, persisting for less than 15 years on the sea floor (Callendar et al., 1994). Red algae and foraminifera dissolved at about 18% per year, which suggests a bottom time of about 6 years. Peak dissolution rates for foraminifera are from 0.02 to 0.09 g/year (Kotler et al., 1992). Aller (1982) noted an overall dissolution rate for shells in Long Island Sound of 1000 – 5000 g CaCO3/m2/year, a great deal more than Keir (1980) found for foraminifera (30 – 800 g CaCO3/m2/year). Powell et al. (1989) listed dissolution capacity (based on titration of all acid produced) for a range of temperate environments. Capacity from 1000 to 4000 g CaCO3/m2/year is common in temperate carbonates from the north Atlantic (both east and west). By comparison, similar calculations in tropical reef environments show dissolution capacity to be only about 800 g CaCO3/m2/ year (Powell et al., 1989). A general dissolution rate for the temperate taphonomically active zone (at and just beneath the sea floor) is about 1000 g CaCO3/m2/ year, nearly twice the carbonate production rate, at least in clastic environments (Davies et al., 1989a). Inorganic carbonate precipitation in tropical environments can be very rapid. For example, Friedman (1998) reported marine cementation of 382 g of oolitic limestone inside a sardine can from a Bahaman tidal flat in only 1 year. As noted earlier, however, sea-floor cementation in temperate regions is rare, and so rates of precipitation are unknown. Once the unique combination of physico-chemical and/or biochemical conditions necessary for marine cementation in coolwater carbonates is achieved, however, cement fabrics indicate (Nelson and James, 2000) that the precipitation process itself could be as rapid as that of tropical

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carbonates (Dravis, 1979; Whittle et al. 1993). For example, Nelson and James (2000) recorded marinecemented shelf-edge temperate carbonate off southern Australia of last post-Glacial age. High-Mg calcite precipitates of inferred Holocene age are reported locally within skeletal chambers from the modern southern Australian shelf (Rahimpour-Bonab et al., 1997) and within intertidal shore platforms developed upon temperate aeolinites along the southeastern Australian coastline (Reeckmann and Gill, 1981). 3.5. Sedimentary implications of dissolution and precipitation In warm water, inorganic precipitation of carbonate is widespread and common. New grains or crystals are precipitated, older grains are infilled, enlarged, or bound together, and structure and strength are added to sedimentary deposits. Dissolution is rare, and localised. This generally constructive regime is in considerable contrast to the situation commonly found in cool- and cold-water carbonates (Table 1). Cooler waters have little or no inorganic precipitation of carbonate, due at least in part to the increased solubility of CaCO3 at lower temperatures. Sea-floor cementation is neither common nor widespread in temperate carbonate deposits (Nelson and James, 2000). On the other hand, oxidation of organic material in surface sediments, enhanced by bioturbation, allows potential for considerable dissolution of metastable carbonate minerals from temperate shelf carbonate sediments, even when overlying sea water is supersaturated with respect to calcite and aragonite. In contrast to physical breakdown, sea-floor dissolution appears to be strongly linked to grain morphology and reactive surface area, as well as the better known factor of mineralogy, so a given skeleton’s susceptibility to dissolution may well be strongly related to its form and thus its taxonomic group. One can readily imagine the preferential dissolution of small delicate bryozoans rather than large robust bivalves in a particular deposit. Smith et al. (1992) compared Recent and Tertiary New Zealand bryozoan faunas and found that erect delicate branching forms which dissolved most readily in the laboratory (Fig. 6) were proportionally more abundant in Recent collections (24% of species) than in Tertiary collections (14%), though of course collection biases and changes

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in environment weaken the comparison. Interestingly, the most dissolution-resistant growth form in their acid-bath experiment was the aragonite bryozoan, Otionella, a relatively dense, compact and free-living form (Fig. 6). In the longer run, mineralogy does have its effect. Early diagenetic dissolution may remove aragonite completely. For example, aragonitic bivalves are entirely missing from many temperate limestones, which are now often dominated by calcitic taxa such as epifaunal bivalves, bryozoans, echinoderms, foraminifera, and barnacles (Nelson, 1978; James and Bone, 1994). Taxonomic bias leads to a ‘‘dissolution fauna’’ where susceptible groups have disappeared because of diagenetic environmental conditions (e.g., Beu et al., 1972). Alternatively, calcitization of aragonite layers in bimineralic bryozoans from the Pleistocene during early diagenesis (e.g., Sandberg, 1975) may allow for preservation of susceptible species. Some young Plio – Pleistocene temperate limestones in New Zealand, where sedimentation rates were relatively high (up to 50 –100 cm/kyr), show a range of preservation, from pristine aragonite shells to chalky aragonite specimens to variably neomorphosed grains to calcite spar-filled biomoulds outlined by micrite envelopes (Hood and Nelson, 1996). This wide variation presumably highlights the importance of microstructural and other species-inherent controls on skeletal stability, in addition to aragonite mineralogy. In Quaternary temperate shelf limestones and aeolinites of coastal southeastern Australia, diagenetic stabilization of aragonite and high-Mg calcite grains has taken from 80 kyr to 1 myr, or longer, during subaerial exposure (Reeckmann and Gill, 1981; Reeckmann, 1988). The most common taxa in temperate deposits (bivalves, bryozoans, echinoderms, foraminifera, and, less commonly, barnacles and red algae) vary in their robustness, microstructure, and mineralogy, and therefore in their susceptibility to dissolution. Bivalves and other large molluscs have robust skeletons (Flessa and Brown, 1983), but they are mainly aragonite. Calcitic bivalves such as scallops and oysters are the most likely to be preserved, and indeed are conspicuous components of most temperate carbonate sediments (e.g., Lees and Buller, 1972; Nelson et al., 1988; Betzler et al., 1997; James 1997). Bryozoans range from robust to very delicate (Smith

et al., 1992), and may be aragonite, calcitic with 0– 10 mol% MgCO3, or of mixed mineralogy (Smith et al., 1998). While most bryozoans are quite susceptible to dissolution, at least one type (robust, compact, and aragonitic) is very resistant (Fig. 6). High-Mg calcitic red algae tend to be very porous and easy to dissolve, while echinoids of a similar mineralogy are rather more resistant to dissolution (Henrich and Wefer, 1986). In practice, incongruent dissolution characterises the diagenetic alteration of many intermediate- to high-Mg calcitic skeletons, leaving behind often wellpreserved low-Mg calcitic replicas (Friedman, 1964; Bathurst, 1975). If dissolution is extensive, carbonate skeletons can fall into their component crystallites (Alexandersson, 1975, 1979). The resulting lime muds are easily transported and may well be flushed from the shelf altogether, or quickly dissolved (Farrow and Fyfe, 1988). Carbonate is thus lost from the shelf environment and any lithified products, perhaps re-entering the global inorganic carbonate cycle. Despite the rarity of temperate marine cements, where they do occur they are responsible for early lithification of cool-water carbonates, preservation of depositional fabrics and of aragonite biomoulds, and even formation of shelf hardgrounds (Nelson and James, 2000). James and Bone (1994) described hardgrounds in mid-Tertiary cool-water limestones in southern Australia which, they suggest, formed whenever the shelf sea floor was brought within the zone of wave abrasion due to fluctuating sea level.

4. Biological sea-floor processes 4.1. Description of biological sea-floor processes In addition to physical and chemical factors, there are biological processes operating on carbonate sediments on the sea floor. Ultimately chemical or physical in nature, biologically mediated processes occupy a spectrum from wholly destructive (bioerosion) to wholly constructive (encrustation), including a wide range of redistribution mechanisms (bioturbation). Bioerosion in skeletal carbonates involves small marine organisms such as algae, fungi, and sponges burrowing, boring, excavating, or otherwise destroying skeletal carbonate, usually for food or protection.

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Many major marine taxa have some boring representatives, usually with atypical lifestyles, life cycles, and life habits. They permanently remove material and/or information from carbonate skeletons, either by boring with acid or by mechanically scraping their way through. Occasionally certain types of bioerosive structures may provide new environmental information, for example, the implication of photic zone residency in shells with blue-green and filamentous green algal microbores (Farrow and Fyfe, 1988). Less commonly, but of some local importance, macrofauna can break or scrape shells. Fish, for example, may fragment the shells of molluscs on which they feed (Cate and Evans 1994). Boring, which can affect the shells of both living and dead organisms (Warme and McHuron, 1978), results in marginal drilling and corrosion of grains, ultimately leading to total breakdown (e.g., Klement and Toomey, 1967). While tropical boring almost always leads to peripheral micritisation of grains, cold-water conditions commonly appear to inhibit the production of micrite (Gunatilaka, 1976). Heavily bored shells tend to be rather more susceptible to abrasive and dissolution forces (Young and Nelson, 1988). Encrustation of shells is a common phenomenon, and often the encrusters are also producers of carbonate. Sessile creatures such as barnacles, bryozoans, attached molluscs, and worms are common encrusters in cool and cold waters. Encrustation is unusual among sea-floor processes, as it is reversible. Any material added to a shell may be knocked off or damaged later, perhaps subsequently to be replaced by new encrusters. There is probably some upper limit of encrustation possible on a given surface, and encrustation may well be seasonal. Encrustation by coralline algae, at least, may reduce potential for bioerosion (Smyth, 1989). Organisms rearrange sediment as they move through it, as they process food, and as they build burrows. Bioturbation is a ubiquitous early sea-floor process that redistributes and eventually homogenises sediment, destroying bedding, usually without adding or subtracting carbonate. In fact, bioturbation occurs irrespective of sediment type and is not specifically related to carbonate budget. In terms of carbonate preservation it tends to be a neutral process, except insofar as it may influence other processes, such as on

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the one hand possibly promoting carbonate dissolution in heavily bioturbated sediment (e.g., Aller, 1982), but on the other also potentially focussing later diagenetic carbonate precipitation within or about the decay traces of organic burrows (e.g., McCunn, 1972). 4.2. Controls on biological sea-floor processes Some carbonate particles are more susceptible to bioerosion than others, even within taxa (Young and Nelson, 1988; Smith, 1993). It remains problematic whether appeal to endoliths reflects ease of access, perhaps mediated by microstructure, or palatability (e.g., more organic layers). Nevertheless, substrate appears to be a primary control (Perry, 1998). Some large aragonite bivalves (Glycymeris and Tucetona) appear particularly susceptible to boring (Young and Nelson, 1988; Farrow and Fyfe, 1988; Gillespie and Nelson, 1997). Massive corals are more extensively bored than free-living or branching corals (Pandolfi and Greenstein, 1997). In contrast, encrustation can occur on any dead substrate. Bioturbation only occurs in unconsolidated sediments and is largely unaffected by sediment texture or composition. The primary extrinsic control on biological processes is assemblage (Table 4). The major microborers are cyanobacteria, green and red algae, and fungi (Pandolfi and Greenstein, 1997). Bioerosion has also been linked to foraminifera and diatoms (Bode´n, 1988). Larger borers and scrapers include bivalves, Table 4 Resistance of various skeletal components to bioerosion (data from Young and Nelson, 1988; Smith, 1993; Pandolfi and Greenstein, 1997)

Taxa

Resistant

Intermediate

Susceptible

Small bivalves (Tawera)

Large gastropods (Struthiolaria)

Large bivalves (Glycymeris, Pecten) Massive corals

Branching corals Texture

Composition

Dense, compact Small

Slightly porous Medium-sized

Highly porous, open Large

Little organic matrix

Average organic matrix

Much organic matrix

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gastropods, chitons, barnacles, bryozoans, and various worms. Even macro-vertebrates play their role, particularly coral-crunching and bottom-feeding fish (Hutchings, 1986). Primary encrusters include worms, algae, barnacles, and bryozoans. Bioturbation is the action of infaunal deposit feeders, including echinoids, holothurians, shrimps, and worms. Assemblage is the result of a combination of environmental factors. There are clear endolithic faunas associated with water depth (and exposure to air in the intertidal range). In the highest intertidal, cyanobacteria are the dominant microborers. At 2– 30 m water depth, red and green algae join the cyanobacteria. At 100 – 300 m depth, fungi become dominant, at least in the Bahamas and Great Barrier Reef settings (Radtke et al., 1996; Vogel et al., 1996). Diversity and abundance also have depth-related distributions; both are greater in the lower subtidal (6– 7 m) than in the upper subtidal (2 –3 m) (Pandolfi and Greenstein, 1997). Similar associations occur with exposure to water energy (Pandolfi and Greenstein, 1997) and general reef setting (Perry, 2000). Encrusters, too, have strong depth-related distributions, similar to those observed on rock walls.

Both encrusters and bioeroders appear to flourish in high-nutrient waters. In highly eutrophic waters, bioerosion may even outstrip carbonate production (Hallock, 1988). The fact that artificial enrichment on the Great Barrier Reef did not result in increased microboring activity may well indicate that endoliths there are not nutrient-limited (Kiene, 1997; Koop et al., 2002). There may be a seasonal influence on these biological processes, at least for bioturbation in temperate environments (e.g., Rice, 1986). Another major control on biological processes is burial. Encrustation does not occur below the sediment – seawater interface. Limited bioerosion occurs to at least 160 cm below the sediment surface, but only cyanobacteria are able to use interstitial nutrients and thrive at such depths (May and Perkins, 1979). Sediment mixing may episodically bury and re-expose shells to the sediment– seawater interface, so both bioerosion and encrustation may be periodic in nature. 4.3. Conceptual models and rates of biological seafloor processes Experimental studies of biologically mediated change in shell weight over time have shown that

Fig. 8. Conceptual model of the effect of biological sea-floor processes of bioerosion and encrustation on skeletal carbonate. Initially encrustation increases carbonate mass, but in the longer term bioerosion outcompetes encrustation with consequently decreased size, loss of grain structure and identity, and production of carbonate mud biodetritus.

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encrustation initially balances or outweighs bioerosion in importance, and shells gain weight (Fig. 8). Over time, however, the increasingly cumulative and irreversible effect of bioerosion leads to a net weight loss and ultimate breakdown of shells. Biologically mediated breakdown occurs in both tropical (Tudhope and Risk, 1985) and temperate (Smith, 1993) waters, but is far more rapid in warm waters. A wide variety of studies have examined rates of bioerosion in reef and other tropical environments. Reefs in Barbados are bioeroded at rates of 100 – 760 g/m2/year, equivalent to 7 –54 cm/kyr (Scoffin et al., 1980). Similarly, reef bioerosion in the Bahamas has been measured at 100– 500 g/m2/year (Vogel, 1997), 250 g/m2/year in Bermuda (Ru¨tzler, 1975), and 350 g/ m2/year on Great Barrier Reef (Tudhope and Risk, 1985). Polychaetes can bioerode up to 10 cm/kyr (Pepe´, 1983), and chitons about 1.7 cm/kyr, or 40 g/ m2/year (Taylor and Way, 1976). Adriatic limestones are bioeroded at 100 cm/year (Torunski, 1979), about 1500 g/m2/year, and Moore and Shedd (1977) report local rates of up to 3 kg/m2/year in warm waters. Average rates of 100– 300 g/m2/year may reach as much as a few kg/m2/year locally or over a short time. There are far fewer studies of temperate bioerosion rates. Farrow and Fyfe (1988) summarise some experimentally determined rates for a variety of shell and limestone substrates in inshore regions from high latitudes, citing typical values from a few to 40 mm/ year. Evans (1968) for example, estimated as much as 12 mm/year erosion on the Oregon coast, mainly brought about by bioerosion. Young and Nelson (1985, 1988) found large temperate bivalves Glycymeris and Humilaria thoroughly bored in 1000 years. Smith (1993) predicted a lifespan of 500 –2000 years for shells on the New Zealand shelf (Fig. 9). Intertidal bioerosion in beachrock of northern California is about 10 – 30 cm/kyr (Stearley and Ekdale, 1989). Rates of encrustation are poorly known. In one study in northern New Zealand molluscs immersed for only 3 months were heavily encrusted; after 12 months worm tubes up to 9 cm long and bryozoans up to 3 cm in diameter were observed (Smith and Nelson, 1994a). Bryozoan production alone was on the order of 0.7 g CaCO3/year, perhaps as much as 24 – 240 g CaCO3/m2/year. An overall sedimentation rate of 4 cm/kyr is consistent with the slow accumulation rates typical of these temperate carbonates (Fig. 3).

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Fig. 9. Short-term changes in weight of various bivalve species immersed in sea water for up to 500 days showing the interplay of encrustation and bioerosion processes (after Smith, 1993).

While each individual infaunal organism redistributes sediment grains in a more or less random way, a whole population operating over a long time can have a fairly predictable effect. Known as the ‘‘eddy biodiffusion coefficient’’ (D) (in cm2/kyr), biodiffusion is related to the concentration of a theoretical ‘‘traceable layer’’ (c), the depth in the sediment (x), and the sedimentation rate (m) (Guinasso and Schink, 1975): D = m(yc/yx)/(y2c/yx2). D varies in shallow water from (10 6 cm2/s) to deep water (10 8 cm2/s). The mixed layer thickness (m) reduces with increasing water depth. The mean thickness of this layer is 9.8 cm over a wide range of environments (Boudreau, 1997), ranging from 5 to 10 cm in the deep sea and up to 1 m in shallow water. A dimensionless parameter ( G) describes the kind of bioturbation acting in a particular environment: G = D/mm. When G is small ( < 0.1), either rapid sedimentation or slow mixing means that burial occurs so rapidly that the effect of bioturbation is small. When G is large (>10) slow sedimentation or rapid mixing cause a great deal of mixing and a nearuniform surface sediment. When these equations are turned around, they can be used to estimate the depth

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at which a particular particle would have been found if bioturbation had not occurred (deconvolution) (Martin, 1999). Such diffusion-based models of bioturbation can only apply if sediment particle mixing is truly random in time and space and on a much smaller scale than the total area under consideration. Selective size feeding, where organisms choose an ideal size range of particles to process, causes a biosorting which makes mixing less like diffusion. Conveyor belt deposit feeders such as shrimps are also less random, and can result in dominantly vertical grain transport. In places where sedimentation is very slow, as is the case for many cool-water carbonate shelves (Fig. 3), mixing can have far more impact than in areas with a high sedimentation rate (Martin, 1999). Rates of bioturbation vary with species. For example, the polychaete worm Pectinaria gouldii ingests 400 ml sediment/worm/year, which amounts to 6000 g/m2/year for the total population, reworking the bed to a depth of 6 cm (Rhoads, 1974). The bivalve Yoldia limatula ingests 257 ml/bivalve/year, 2262 g/m2/year for the total population, and the depth of reworking is 0– 2 cm (Rhoads, 1974). However, such rates are for species living in muddy or sandy sediments, and may not be applicable to the coarser carbonate hashes typical of most temperate shelves. No study of reworking in such a setting is known to us. 4.4. Sedimentary implications of biological sea-floor processes Bioerosion has been invoked as a major destructive process in both tropical and temperate waters (e.g., Neumann, 1966; Perkins and Halsey, 1971; Vogel et al., 1999). Grains are first pitted, bored, and etched by bioerosion, and in the later stages, exfoliation and extreme fragility can result. Shells that have been subject to bioerosion have increased susceptibility to mechanical stress (Young and Nelson, 1988); even a single drill-hole can reduce strength. The end result is destruction of information, and reduction in grain size. The carbonate mud produced by abrasive borers is often distinctive. Clionid sponges, in particular, produce characteristic 15- to 100-Am chips, which may comprise up to 30% of total sediment in some tropical environments (Fu¨tterer, 1974).

Encrusted grains may (e.g., Silva de Echols, 1993) or may not (Chave, 1964) be strengthened and protected by the biogenic carbonate precipitated around them (Smyth, 1989). Encrusters add new carbonate to the system increasing grain size and overall solid carbonate present, and sometimes binding grains together, even perhaps forming small biogenic structures. Over time, skeletons may acquire their own stratigraphy, with an original substrate increasingly obscured by subsequent layers of carbonate (e.g., Gherardi and Bosence, 1999). Encrusters provide information about the environment in which the original skeleton has rested and been colonised. It is a constructive process, although periodic and possibly seasonal. It is common to characterise sediment grains as ‘‘relict’’ or ‘‘older’’ by examination of surface features. Bioerosion and encrustation are obvious evidence of time spent on the sea floor, and may well be the primary features used in visual determinations of relative age (but see, e.g., Powell and Davies, 1990). Carbon dating, however, shows that there is considerable variation among species, and that relatively unbored and unencrusted grains of a resistant or unappealing species may be older than decrepit-looking shells of an attractive or susceptible species (Nelson and Bornhold, 1983; Nelson et al., 1988; Smith 1993). If large molluscs can be destroyed in 2000 years or less (Smith, 1993; Gillespie and Nelson, 1997), then quite degraded-looking shells could be modern and not necessarily relict. Nevertheless, some paleoenvironmental inferences can be based on bioeroders and encrusters. The presence or absence of photosynthetic borers and encrusters, for example, gives a clear indication of paleodepth within the photic zone (e.g., Klement and Toomey, 1967; Perkins and Halsey, 1971). Similarly, different reef environments appear to be characterised by distinct boring faunas, allowing reconstruction of reef deposits (Vogel et al., 1999; Perry, 2000). Bioeroded hardgrounds associated with carbonate depositional hiatuses may indicate periods where carbonate production was outstripped by bioerosion, perhaps due to high nutrient availability (Hallock, 1988). Although bioerosion has been documented since the Pre-Cambrian, it is likely that the intensity of attack on carbonate may not have been so great in the past. Clionid sponges, for example, did not become abun-

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dant in the fossil record until the Triassic, and excavating grazers were unimportant until the Cretaceous, and scraping fishes only appeared in the Miocene (Vogel, 1993). Bioturbation may move skeletal grains to and from the sediment/seawater interface, which affects their degree of exposure to sea-floor and near-below seafloor diagenetic processes. It has many geochemical implications, including the deepening of the oxidised zone, the early diagenesis of sulfur and carbon (Berner and Westrich, 1985), and sediment porosity and permeability. Temperate bioerosion appears to be slow enough to require long residence times at the sediment – sea water interface for significant destruction to occur, on the order of hundreds to a few thousands of years for total shell destruction (Smith, 1993). Sedimentation rates on temperate shelves are often sufficiently low for both bioerosion and encrustation to have significant effects. Both boring and encrusting will cease when a shell is buried, and begin again when it

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is re-exposed. Bioturbation is ubiquitous and consistently important.

5. Concluding remarks 5.1. The temperate early sea-floor regime and budget The contributions made by physical, chemical, and biological early sea-floor processes to either adding or destroying information about an accumulating shelf carbonate deposit are summarised in Fig. 10. It has long been known that cool-water carbonate deposits are subjected to dominantly destructive early marine diagenesis (e.g., Nelson, 1988a; James, 1997). While production rates may be locally similar to tropical environments (up to 2000 g CaCO3/m2/year), longterm accumulation of temperate carbonates is far slower (Fig. 3). The difference is made up by the destructive nature of the early sea-floor processes characterising temperate shelves (Smith, 1988). A

Fig. 10. Constructive and destructive nature of early sea-floor processes operating in shelf carbonate sediments. In cool-water settings, destructive conditions typically outweigh constructive ones, in many situations apparently entirely so, producing characteristic cool-water shelf carbonate attributes (Tables 1, 6 and 7).

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temperate carbonate sediment is typically uncemented, abraded, bored and without micrite infillings or inorganic grains (see, e.g., Nelson and Bornhold, 1984). While dissolution may be locally or episodically important (accounting for 1000 –5000 g CaCO3/m2/ year), most temperate carbonates are more influenced by bioerosion. Boring may destroy some 100– 300 g CaCO3/m2/year, but as it is continuous and ubiquitous, is probably of more importance worldwide. While abrasion on temperate shelves is probably less than that measured in the surf zone (100 – 400 g CaCO3/m2/ year), the influence of hard siliciclastic material admixed with the skeletal hashes may make it quite significant, particularly on the inner shelf. Outer shelf environments may experience little or no abrasion (e.g., James and Bone, 1991), except under fluctuating sealevel conditions (e.g., Boreen et al., 1993). These same destructive influences operate in shallow tropical seas, but there they are balanced by a series of constructive influences, most importantly inorganic precipitation of carbonate. Cementation, precipitation, and higher production associated with coral reefs lead to rapid burial and escape from the forces of bioerosion and abrasion. A rough annual carbonate budget for a typical reef (based on figures from Powell et al., 1989) and a typical temperate shelf (Table 5) shows that preservation in the reef environment may be about 50– 75% (and can be higher depending on production), whereas it may only reach 25% in the temperate shelf. If production is low or dissolution active, carbonate preservation on temperate shelves could be close to

Table 5 Comparison of carbonate budgets from a typical coral reef and a typical temperate shelf carbonate deposit (based on figures from this paper; Smith, 1988; Powell et al., 1989; Wollast, 1994) Coral reef (in g CaCO3/ m2/year)

Temperate shelf carbonate (in g CaCO3/ m2/year)

Production and precipitation and cementation

5000

1600

Dissolution Bioerosion Abrasion

800 350 50 – 1500

1000 – 5000 100 – 350 100 – 400

Net accumulation

2350 – 3800

0 – 400

Preservation rate

47 – 76%

0 – 25%

zero. Many of the figures used in this calculation are ‘‘best-guess’’ estimates and a better understanding of budgets depends on acquiring more accurate and more widespread data from temperate regimes, as noted by Milliman and Droxler (1996). The considerable variability in the effects of early sea-floor processes on temperate carbonate sediments is illustrated in Table 5. Wide ranges of rates reflect variability among sedimentary environments. There are two major controls on destructive diagenesis: assemblage and sedimentation. 5.2. Effects of assemblage Temperate carbonate deposits are made up entirely of skeletons of organisms. A ‘‘natural archive’’ (Powell et al., 1989, p. 555), they contain a record of environment and ecology of the past. While the archive is flawed by the absence of non-preservable organisms, it is still commonly used to elucidate paleoenvironment. Presence/absence data may be fairly accurate (among skeleton-forming taxa), but abundance and importance may be strongly affected by diagenesis, particularly the early sea-floor processes which act selectively (Fig. 11). Different taxa are more or less resistant to the destructive temperate regime (Tables 2, 3 and 4). The Sorby Principle may well be applicable to the major temperate taxa. Intrinsic characteristics of the skeletons (composition, microstructure, shape, and size) eventually determine their response to abrasion, dissolution, and bioerosion. It has been demonstrated that abrasion and dissolution, in particular, act differently on different gross morphologies, relating to compactness and reactive surface area (e.g., Smith and Nelson, 1994b). Dissolution may act more rapidly on aragonitic grains than on calcitic ones. The observation that some species of mollusc are more susceptible to bioerosion begs the question—what characteristics of the shell influence bioeroders’ choices? Temperate carbonates are typically dominated by molluscs, particularly bivalves. Other dominant taxa include bryozoans and foraminifera, with echinoids, barnacles, and coralline rhodophytes making significant contributions in some areas. While significant variation occurs within taxa, some generalisations may be made (Hayton et al., 1995). Large compact

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Fig. 11. An example of the effects of early sea-floor processes on a cool-water deposit (after Smith and Nelson, 1994b). On southern temperate shelves, bryozoans are a dominant component; this model begins with typical bryozoan assemblages, applying the effects of physical, chemical, and biological processes, and predicting the eventual taphofacies available for preservation. Bryozoan growth forms: ENml—encrusting multilaminar; ENul—encrusting unilaminar; ERro—erect rigid robust branching; ERde—erect rigid delicate branching; ERfe—erect rigid fenestrate; ERfo—erect rigid foliose; EFar—erect flexible articulated; and FLmo—free-living motile (following scheme of Nelson et al., 1988).

bivalves are most resistant to abrasion and early dissolution, but they are often particularly attractive to bioeroders, and if they are aragonite, they can be altered or removed in later diagenesis. Most bryozoans and calcareous red algae are quite fragile and can be easily disarticulated; a few are more robust. Barnacles, echinoids, and algae (are both HMC)are mostly calcitic, and the high-Mg calcite of coralline rhodophytes is

particularly susceptible to dissolution or alteration (though rhodoliths are often well-preserved). Check intention of last sentence. 5.3. Effects of sedimentation rate One of the key assumptions in taphonomy is that preservation is enhanced by rapid burial, effectively

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ending the regime of early sea-floor processes; this is the ‘‘Second Great Rule of Taphonomy’’ of Martin (1999). The obvious corollary is that long residence time at the sediment/seawater interface (i.e., in the taphonomically active zone) results in greater effects of sea-floor processes. In the destructive diagenetic regime of temperate shelves, a result of long-term exposure will be decreasing grain size accompanied by information loss. Clearly, burial will stop the effects of abrasion and of encrustation. Dissolution and cementation may persist further down into the sediment depending on pore-water chemistry. It has been shown that bioerosion can take place up to 160 cm below the sediment/seawater interface (May and Perkins, 1979). Nevertheless, long residence time at the sea bed allows all three major destructive processes to act on sediment. Bioerosion sometimes begins before death (Gillespie and Nelson, 1997) and could continue after lithification (Martin, 1999) such that only removal from the normal marine environment, either by burial or uplift, can stop it. Some large aragonitic bivalves may be completely bioeroded (leaving out dissolution and abrasion) in about 1000 years (Smith, 1993). If most bioerosion occurs in the taphonomically active zone (say the upper 20 cm of sediment), then sedimentation rate must be greater than 20 cm/kyr in order for these large bivalves to be preserved. In fact, temperate carbonate shelves have far lower sedimentation rates (Table 1). In contrast with tropical reefs and carbonate environments, where sedimentation ranges from about 3.5 to 1400 cm/kyr (average of 18 samples = 250 cm/ kyr), and coastal wetlands where it can be as much as 8000 cm/kyr (see summary in Martin, 1999, p. 280), temperate shelves typically exhibit overall sedimentation rates less than 10 cm/kyr (Nelson, 1988b; James, 1997). Such slow sedimentation rates not only enhance effectiveness of destructive early diagenetic processes, but also result in greater timeaveraging of assemblages and deposits. We can imagine a given sedimentary layer being buried. At first, it accumulates at the sediment/seawater interface, and is subjected to early sea-floor processes. Over time it is buried, reworked by bioturbation, re-exposed, and re-buried until it reaches ‘‘maturity’’ (sensu Martin, 1999), passes into the

historical record, and is no longer subject to reworking. This thoroughly mixed sediment becomes subject to ‘‘true’’ diagenesis and eventually may be lithified and preserved. In low sedimentation regimes such as the temperate carbonate shelf, the mixed sediment layer may be quite thin (5– 10 cm) and yet represent time-averaging of hundreds to thousands of years (Smith, 1993; Martin, 1999). In most cases, carbonate skeletons last longer on the shelf than they do in experiments. This is probably due to episodic burial and re-exposure. Even a short period of burial reduces the effects of taphonomic processes (Cummins et al., 1986; Powell et al., 1989). Among the main processes of burial, net sedimentation and catastrophic burial are not so important on the temperate shelf as bioturbation and associated resuspension and redeposition. 5.4. Cumulative effects Early sea-floor processes do not operate in isolation or in series, but in tandem. The cumulative effects of destructive sea-floor processes may be synergistic, as each process appears to increase vulnerability to others (e.g., Young and Nelson, 1988). It is the interaction of all processes acting on a grain or a sediment deposit that determine preservation potential and final form. Boring and burrowing, in particular, weaken grains and enhance other destructive processes. Bioerosion increases reactive surface area, and may well allow later diagenetic chemical processes greater scope for change. Certainly, bored shells degrade to carbonate mud much faster than fresh ones (Young and Nelson, 1988). Unlike in tropical carbonate environments, boring holes in temperate shelf settings are not commonly filled with micrite, so they remain potential weakening agents for the life of the particle (Young and Nelson, 1988). Abrasion, too, enhances reactive surface area, and may expose new surfaces for bioeroders to colonise. Similarly, constructive processes enhance preservation of both material and information. Encrustation and cementation protect their substrate in a kind of ‘‘micro-burial’’ away from destructive processes. Although some encrusters are borers, and remove shell material as they cover it, most act to preserve their host material.

Table 6 Summary of early sea-floor processes in temperate shelf carbonate sediments Dissolution

Cementation

Bioerosion

Encrustation

Bioturbation

Description

Mechanical breakdown

Chemical solution

Chemical precipitation

Boring by organisms

Overgrowth by organisms

Disturbance of sediment by organisms

Symptoms

Rounded, smooth grains

Pitted, etched grains

Crusts, rinds, infilling

Pits, bores, etching

Skeletal overgrowths

Mixed sediments

Intrinsic controls

Composition Size and shape Microarchitecture

Reactive surface area Mineralogy

Carbonate available

Appeal to endoliths ?Organic matrix

Hard substrate available

Extrinsic controls

Hydraulic regime Terrigenous sediment mixing Exposure time

Saturation state of water Bioturbation Exposure time

Saturation state of water Hydraulic regime Exposure time

Taxa present Environment especially nutrients Exposure time

Taxa present Environment Hydraulic regime Exposure time

Sedimentation rate Taxa present

Model

Wt = W0e at where a = f (E,I) (see Fig. 4)

R = k(1 X)n (see Fig. 7)

Wt = W0 + f (X,t) (see Fig. 7)

see Fig. 9

see Fig. 9

G = D/mm (see text)

Rates (g CaCO3/m2/ year) (percentage)

100 – 400, 12%/km surf zone transport

about 1000 18% per year

Unknown

100 – 350, 10 – 20% per year

Unknown

1000 – 6000 g/m2/ year reworking

Sedimentological implications

Destruction of information Reduction in size Increase in sorting Increased susceptibility to dissolution Source of carbonate mud

Destruction of information Reduction in size Source of carbonate mud

Clouds biological information Adds diagenetic information Adds carbonate Reduces susceptibility to destructive processes Lithification

Destruction of information Reduction in strength, size Adds some environmental information Increases susceptibility to dissolution Source of carbonate mud

Preservation of carbonate Increase in strength Adds post-depositional information

May periodically bury and re-expose sediments Mixing of time horizons

Typical fossils or deposits

Large compact molluscs Robust bryozoans Benthic forams

Large calcitic molluscs Robust calcitic bryozoans Barnacles

Mainly calcitic cements Often intermediateMg calcite

Most taxa have a boring representative

Most sessile marine taxa are can be encrusting

Most marine sediments are bioturbated

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Abrasion

Model variables: W0 = original weight of particle; Wt = weight of particle after time t; a = constant rate of abrasion; t = time; E = extrinsic factors; I = intrinsic factors; R = dissolution rate; k = dissolution rate constant (reactive surface area); X = saturation state of water; n = order of reaction (depends on mineralogy); G = bioturbation parameter; D = eddy biodiffusion coefficient; m = mixed layer thickness; m = sedimentation rate. 25

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Table 7 The temperate shelf carbonate taphofacies associated with destructive early sea-floor processes (based on data from this paper and using criteria adapted from Powell et al., 1989) Parameters

Temperate shelf carbonate characteristics

Attributes of skeletal assemblage Taxa present Molluscs, bryozoans, barnacles, echinoids, foraminifera, coralline red algae Taxa absent Hermatypic corals, codiacean green algae Wholeness Generally broken down into fragments, disarticulated, periostracum missing Orientation Not in life position, usually random orientations Attributes of skeletal surfaces Abrasion features Major abrasion evident in fragmentation, chipped and rounded edges, smoothing of architecture, overall rounding Dissolution features Minor dissolution, if any, resulting in chalkiness, some pitting and dulling of surface features Cementation features Usually none Bioerosion features Borings and pits common and extensive Encrustation features Encrusting common

Parameters for taphofacies have been usefully summarised by Powell et al. (1989), and here, we apply some of them to temperate shelf carbonates (Table 7). The overall temperate ‘‘bryomol’’ to ‘‘foramol’’ assemblage, taken together with the lack of inorganic precipitation, points to a cool-water origin for the deposit. Subsequent long exposure to abrasion and biologically mediated processes (and possibly episodic dissolution) results in rounded, smoothed, and sometimes chalky fragments. On some temperate shelves, it may be that destructive diagenesis is sufficient to account for all production, so that no carbonate is preserved in buried sediment or the fossil record. In this destructive regime, even substantial production of carbonate could be removed and therefore remain poorly known or unrecorded. It may well be that carbonates of the temperate shelf were common and widespread in the past, as they are today, but long-term exposure to destructive diagenetic processes at the sea floor substantially reduced their preservation potential, possibly leading to a bias toward tropical limestones in the rock record.

5.5. The temperate shelf carbonate taphofacies

Acknowledgements

On temperate shelves the taphonomically active zone is a lively place. The upper few to tens of centimeters (Powell et al., 1989) are host to chemical, physical, and biological processes, all of which alter skeletal carbonate. The most active processes are destructive, and each of them is selective—they act differently on different sediment grains (e.g., Fig. 11). Characteristics of both the grain and the environment influence the effects of early sea-floor processes, and how selection works (Table 6). The effect of selective destruction is to bias the sedimentary record and render it inaccurate in terms of the assemblage and environment it represents. Since both the extent of preservation and the degree of bias are unknown, it would seem that any interpretation of the sedimentary record ought to be deeply flawed, and that much of the blame rests with early-sea floor processes. Nevertheless, these processes themselves may leave evidence behind, making the final deposit a record of both the original environment and the many processes which have acted on it—a taphofacies (e.g., Fig. 11).

This work was supported by the New Zealand University Grants Committee and the University of Waikato Research Committee Grant RF-SF-S487, and by a University of Waikato Postgraduate Scholarship to AMS. We thank Oliver Gussmann (University of Otago) for his constructive comments on the manuscript, and John Milliman (Virginia Institute of Marine Science, College of William and Mary) and Murray Gregory (University of Auckland) for their useful reviews. References Aguirre, M.L., Farinati, E.A., 1999. Taphonomic processes affecting late Quaternary molluscs along the coastal area of Buenos Aires Province (Argentina, southwestern Atlantic). Palaeogeography, Palaeoclimatology, Palaeoecology 149, 283 – 304. Alexandersson, T., 1972. Shallow-marine carbonate diagenesis as related to the carbonate saturation level in seawater. Publications from the Palaeontological Institution of the University of Uppsala 126, 1 – 10. Alexandersson, T., 1974. Carbonate cementation in coralline algal nodules in the Skagerrak, North Sea: biochemical precipitation

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Zealand Cenozoic limestones. Sedimentary Geology 100, 123 – 141. Henrich, R., Wefer, G., 1986. Dissolution of biogenic carbonates: effects of skeletal structure. Marine Geology 71, 341 – 362. Hood, S.D., Nelson, C.S., 1996. Cementation scenarios for New Zealand Cenozoic nontropical limestones, New Zealand. Journal of Geology and Geophysics 39, 109 – 122. Hoskin, C.M., Geier, J.C., Reed, J.K., 1983. Sediment produced from abrasion of the branching stony coral Oculina varicosa. Journal of Sedimentary Petrology 53, 779 – 786. Hutchings, P.A., 1986. Biological destruction of coral reefs: a review. Coral Reefs 4, 239 – 252. Irwin, H., Curtis, C.D., Coleman, M., 1977. Isotopic evidence for source of diagenetic carbonates in organic-rich sediments. Nature 269, 209 – 213. James, N.P., 1997. The cool-water carbonate depositional realm. In: James, N.P., Clarke, J.A.D. (Eds.), Cool-water Carbonates, vol. 56. SEPM Special Publication, Tulsa, OK, pp. 1 – 20. James, N.P., Bone, Y., 1991. Origin of a cool-water Oligo – Miocene deep-shelf limestone, Eucla Platform, Southern Australia. Sedimentology 38, 323 – 341. James, N.P., Bone, Y., 1992. Synsedimentary cemented calcarenite layers in Oligo – Miocene shelf limestones, Eucla Platform, southern Australia. Journal of Sedimentary Petrology 62, 860 – 872. James, N.P., Bone, Y., 1994. Paleoecology of cool-water, subtidal cycles in mid-Cenozoic limestones, Eucla Platform, South Australia. Palaios 9, 457 – 476. James, N.P., Clarke, J.A.D. (Eds.), 1997. Cool-water Carbonates, vol. 56. SEPM Special Publication, Tulsa, OK. Jørgensen, N.O., 1976. Recent high magnesian calcite/aragonite cementation of beach and submarine sediments from Denmark. Journal of Sedimentary Petrology 46, 940 – 951. Keir, R.S., 1980. The dissolution kinetics of biogenic calcium carbonates in seawater. Geochimica et Cosmochimica Acta 44, 241 – 252. Kidwell, S.M., Baumiller, T., 1990. Experimental disintegration of regular echinoids: roles of temperature, oxygen, and decay thresholds. Paleobiology 16, 247 – 271. Kiene, W.E., 1997. Enriched nutrients and their impact on bioerosion: results from ENCORE. Proceedings 8th International Coral Reef Symposium, vol. 1, pp. 897 – 902. Klement, K.W., Toomey, D.T., 1967. Role of the blue-green alga Girvanella in skeletal grain destruction and lime-mud formation in the lower Ordovician of west Texas. Journal of Sedimentary Petrology 37, 1045 – 1051. Koop, K., Brodie, J., Bucher, D., Capone, D., Coll, J., Dennison, W., Erdmann, M., Harrison, P., Hoegh-Guldberg, O., Hutchings, P., Jones, G.B., Larkum, A.W.D., O’Neil, J., Steven, A., Tentori, E., Ward, S., Williamson, J., Yellowlees, D., Booth, D., Broadbent, A., 2002. ENCORE: The effect of nutrient enrichment on coral reefs. Synthesis of results and conclusions. Marine Pollution Bulletin 42, 91 – 120. Kotler, E., Martin, R.E., Liddell, W.D., 1992. Experimental analysis of abrasion and dissolution resistance of modern reef-dwelling foraminifera: implications for the preservation of biogenic carbonate. Palaios 7, 244 – 276.

A.M. Smith, C.S. Nelson / Earth-Science Reviews 63 (2003) 1–31 Kukal, Z., 1990. The rate of geological processes. Earth-Science Reviews 28, 1 – 259. Lees, A., Buller, A.T., 1972. Modern temperate-water and warmwater shelf carbonate sediments contrasted. Marine Geology 13, M67 – M73. Lewy, Z., 1975. Early diagenesis of calcareous skeletons in the Baltic Sea, western Germany. Meyniana 27, 29 – 33. Lewy, Z., 1981. Maceration of calcareous skeletons. Sedimentology 28, 893 – 895. Maiklem, W.R., 1968. Some hydraulic properties of bioclastic grains. Sedimentology 10, 101 – 109. Martin, R.E., 1999. Taphonomy: A Process Approach. Cambridge Paleobiology Series, vol. 4. Cambridge Univ. Press, Cambridge. 508 pp. Matthews, E.R., 1983. Measurements of beach pebble attrition in Palliser Bay, southern North Island, New Zealand. Sedimentology 30, 787 – 799. May, J.A., Perkins, R.D., 1979. Endolithic infestation of carbonate substrates below the sediment – water interface. Journal of Sedimentary Petrology 49, 357 – 378. McCunn, H.J., 1972. Calcite and aragonite phenomena precipitated by organic decay in high lime concentrate brines. Journal of Sedimentary Petrology 42, 150 – 154. Menard, H.W., Boucot, A.J., 1951. Experiments on the movement of shells by water. American Journal of Science 249, 131 – 151. Milliman, J.D., Droxler, A.W., 1996. Neritic and pelagic carbonate sedimentation in the marine environment: ignorance is not bliss. International Journal of Earth Sciences (Geologische Rundschau) 85, 496 – 504. Mitchell-Tapping, H.J., 1981. Particle breakdown of recent carbonate sediment in coral reefs. Florida Scientist 44, 21 – 29. Moberly Jr., R., 1970. Microprobe study of diagenesis in calcareous algae. Sedimentology 14, 113 – 123. Moore, C.H., Shedd, W.W., 1977. Effective rates of sponge bioerosion as a function of carbonate production. Proceedings of the Third International Coral Reef Symposium, vol. 2, pp. 499 – 505. Morse, J.W., Mackenzie, F.T., 1990. Geochemistry of Sedimentary Carbonates. Developments in Sedimentology, vol. 48. Elsevier, Amsterdam. 707 pp. Morse, J.W., Mucci, A., Millero, F.J., 1980. The solubility of calcite and aragointe in seawater of 35 ppt salinity at 25 jC and atmospheric pressure. Geochimica Cosmochimica Acta 44, 85 – 94. Nelson, C.S., 1978. Temperate shelf carbonate sediments in the Cenozoic of New Zealand. Sedimentology 25, 737 – 771. Nelson, C.S. (Ed.), 1988a. Non-tropical Shelf Sediments—Modern and Ancient. Sedimentary Geology, vol. 60, pp. 1 – 367. Nelson, C.S., 1988b. An introductory perspective on non-tropical shelf carbonates. Sedimentary Geology 60, 3 – 12. Nelson, C.S., Bornhold, B.D., 1983. Temperate skeletal carbonate sediments on Scott shelf, northwestern Vancouver Island, Canada. Marine Geology 52, 241 – 266. Nelson, C.S., Bornhold, B.D., 1984. Temperate continental shelf skeletal carbonate deposits. Geobios Memoir Special 8, 109 – 113. Nelson, C.S., Hancock, G.E., 1984. Composition and origin of temperate skeletal carbonate sediments on South Maria Ridge, northern New Zealand. New Zealand Journal of Marine and Freshwater Research 18, 221 – 239.

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Nelson, C.S., James, N.P., 2000. Marine cements in mid-Tertiary cool-water shelf limestones of New Zealand and southern Australia. Sedimentology 47, 609 – 629. Nelson, C.S., Lawrence, M.F., 1984. Methane-derived high-Mg calcite submarine cement in Holocene nodules from the Fraer Delta, British Columbia, Canada. Sedimentology 31, 645 – 654. Nelson, C.S., Hancock, G.E., Kamp, P.J.J., 1982. Shelf to basin, temperate skeletal carbonate sediments, Three Kings Plateau, New Zealand. Journal of Sedimentary Petrology 52, 717 – 732. Nelson, C.S., Keane, S.L., Head, P.S., 1988. Non-tropical carbonate deposits on the modern New Zealand shelf. Sedimentary Geology 60, 71 – 94. Neumann, A.C., 1966. Observations on coastal erosion in Bermuda and measurements of the boring rate of the sponge, Cliona lampa. Limnology and Oceanography 11, 92 – 108. Nicolaides, S., Wallace, M.W., 1997. Pressure-dissolution and cementation in an Oligo – Miocene non-tropical limestone (Clifton Formation), Otway Basin, Australia. In: James, N.P., Clarke, J.A.D. (Eds.), Cool-water Carbonates, vol. 56. SEPM Special Publication, Tulsa, OK, pp. 249 – 261. Pandolfi, J.M., Greenstein, B.J., 1997. Taphonomic alteration of reef corals: effects of reef environment and coral growth form: I. The Great Barrier Reef. Palaios 12, 27 – 42. Parsons-Hubbard, K.M., Callender, W.R., Powell, E.N., Brett, C.E., Walker, S.E., Raymond, A.L., Staff, G.M., 1999. Rates of burial and disturbance of experimentally deployed molluscs: implications for preservation potential. Palaios 14, 337 – 351. Pepe´, P.J., 1983. Bioerosion by a polychaete annelid, Eunice afra Peters, at Puerto Pan˜asco, Gulf of California, Mexico. Unpublished PhD thesis, University of Southern California, Los Angeles, 179 pp. Perkins, R.D., Halsey, S.D., 1971. Geologic significance of microboring fungi and algae in Carolina shelf sediments. Journal of Sedimentary Petrology 41, 843 – 853. Perry, C.T., 1998. Grain susceptibility to the effects of microboring: implications for the preservation of skeletal carbonates. Sedimentology 45, 39 – 51. Perry, C.T., 2000. Macroboring of Pleistocene coral communities, Falmouth Formation, Jamaica. Palaios 15, 483 – 491. Pilkey, O.H., Fierman, E.I., Trumbull, J.V.A., 1979. Relationship between physical condition of the carbonate fraction and sediment environments: northern Puerto Rico shelf. Sedimentary Geology 24, 283 – 290. Powell, E.N., Davies, D.J., 1990. When is an ‘‘old’’ shell really old? Journal of Geology 98, 823 – 844. Powell, E.N., Staff, G.M., Davies, D.J., Callender, W.R., 1989. Macrobenthic death assemblages in modern marine environments: formation, interpretation, and application. Reviews in Aquatic Sciences 1, 555 – 589. Pytkowicz, R.M., 1965. Calcium carbonate saturation in the oceans. Limnology and Oceanography 10L, 220 – 225. Radtke, G., Le Campion-Slumard, T., Stjepko, G., 1996. Microbial assemblages of the bioerosional ‘‘notch’’ along tropical limestone coasts. Algological Studies 83, 469 – 482. Rahimpour-Bonab, H., Bone, Y., Moussavi-Harami, R., Turnbull, K., 1997. Geochemical comparisons of modern cool-water calcareous biota, Lacepede Shelf, South Australia, with their tropical

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counterparts. In: James, N.P., Clarke, J.A.D. (Eds.), Cool-water Carbonates, vol. 56. SEPM Special Publication, Tulsa, OK, pp. 77 – 91. Rao, C.P., 1981. Cementation in cold-water bryozoan sand, Tasmania, Australia. Marine Geology 40, M23 – M33. Rao, C.P., 1996. Modern Carbonates—Tropical, Temperate, Polar. Carbonates, Tasmania, 206 pp. Reeckmann, S.A., 1988. Diagenetic alterations in temperate shelf carbonates from southeastern Australia. Sedimentary Geology 60, 209 – 219. Reeckmann, S.A., Gill, E.D., 1981. Rates of vadose diagenesis in Quaternary dune and shallow marine calcarenites, Warrnambool, Victoria, Australia. Sedimentary Geology 30, 157 – 172. Rhoads, D.C., 1974. Organism – sediment relations on the muddy sea floor. Oceanographic Marine Biology Annual Review 12, 263 – 300. Rice, D.L., 1986. Early diagenesis in bioadvective sediments: relationships between the diagenesis of beryllium-7, sediment reworking rates, and the abundance of conveyor-belt depositfeeders. Journal of Marine Research 44, 149 – 184. Ru¨tzler, K., 1975. The role of burrowing sponges in bioerosion. Oecologia 19, 203 – 216. Sandberg, P.A., 1975. Bryozoan diagenesis: bearing on the nature of the original skeleton of rugose corals. Journal of Paleontology 49, 587 – 604. Scoffin, T.P., Stearn, C.W., Boucher, D., Frydl, P., Hawkins, C.M., Hunter, I.G., MacGeachy, J.K., 1980. Calcium carbonate budget of a fringing reef on the west coast of Barbados. Part II—Erosion, sediments and internal structure. Bulletin of Marine Science 30, 475 – 508. Shroba, C.S., 1993. Taphonomic features of benthic Foraminifera in a temperate setting: experimental and field observation on the role of abrasion, solution and microboring in the destruction of foraminiferal tests. Palaios 8, 250 – 266. Silva de Echols, C.M.H.M., 1993. Diatom infestation of recent crinoid ossicles in temperate waters, Friday Harbor Laboratories, Washington: implications for biodegradation of skeletal carbonates. Palaios 8, 278 – 288. Smith, A.M., 1988. Preliminary steps toward formation of a generalized budget for cold-water carbonates. Sedimentary Geology 60, 323 – 331. Smith, A.M., 1993. Bioerosion of bivalve shells in Hauraki Gulf, North Island, New Zealand. In: Battershill, C.N., et al., (Eds.), Proceedings of the Second International Temperate Reef Symposium, Auckland, New Zealand. NIWA Marine, Wellington, pp. 175 – 181. Smith, A.M., Nelson, C.S., 1994a. Calcification rates of rapidlycolonising bryozoans in Hauraki Gulf, northern New Zealand. New Zealand Journal of Marine and Freshwater Research 28, 227 – 234. Smith, A.M., Nelson, C.S., 1994b. Selectivity in sea-floor processes: taphonomy of bryozoans. In: Hayward, P.J., Ryland, J.S., Taylor, P.D. (Eds.), Biology and Palaeobiology of Bryozoans. Olsen and Olsen, Fredensborg, pp. 177 – 180. Smith, A.M., Nelson, C.S., 1996. Differential abrasion of bryozoan skeletons: taphonomic implications for paleoenvironmental interpretation. In: Gordon, D.P, Smith, A.M., Grant-Mackie, J.A.

(Eds.), Bryozoans in Space and Time, Proceedings of the 10th International Bryozoology Conference, Wellington, New Zealand, 1995. National Institute of Water and Atmospheric Research, Wellington, pp. 305 – 313. Smith, A.M., Nelson, C.S., Danaher, P.J., 1992. Dissolution behaviour of bryozoan sediments: taphonomic implications for nontropical shelf carbonates. Palaeogeography, Palaeoclimatology, Palaeoecology 93, 213 – 226. Smith, A.M., Nelson, C.S., Spencer, H.G., 1998. Skeletal mineralogy of New Zealand bryozoans. Marine Geology 151, 27 – 46. Smith, S.V., 1971. Budget of calcium carbonate, Southern California continental borderland. Journal of Sedimentary Petrology 41, 798 – 808. Smyth, M.J., 1989. Bioerosion of gastropod shells: with emphasis on effects of coralline algal cover and shell microstructure. Coral Reefs 8, 119 – 125. Sneed, E.D., Folk, R.L., 1958. Pebbles in the Lower Colorado River, Texas. A study in particle morphogenesis. Journal of Geology 66, 114 – 150. Sorby, H.C., 1879. The structure and origin of limestones. Geological Society of London Proceedings 35, 56 – 95. Stearley, R.F., Ekdale, A.A., 1989. Modern marine bioerosion by macroinvertebrates, northern Gulf of California. Palaios 4, 453 – 467. Taylor, J.D., Way, K., 1976. Erosive activities of chitons at Aldabra Atoll. Journal of Sedimentary Petrology 46, 974 – 977. Trimonis, E.S., 1974. Some characteristics of carbonate sedimentation in the Black Sea. In: Degas, E.T., Ross, P.A. (Eds.), The Black Sea—Geology, Chemistry and Biology. American Associations for Petroleum Geologists Memoir, Tulsa, OK, pp. 279 – 295. Torunski, H., 1979. Biological erosion and its significance for the morphogenesis of limestone coasts and for nearshore sedimentation (Northern Adriatic). Senckenbergiana Maritima 11, 193 – 265. Tucker, M.E., Bathurst, R.C.C., 1990. Carbonate Diagenesis. IAS Reprint Series, vol. 1. Blackwell, Oxford, 312 pp. Tucker, M.E., Wright, V.P., 1990. Carbonate Sedimentology. Blackwell, Oxford. 482 pp. Tudhope, A.W., Risk, M.J., 1985. Rate of dissolution of carbonate sediments by microboring organisms, Davies Reef, Australia. Journal of Sedimentary Petrology 55, 440 – 447. Vogel, K., 1993. Bioeroders in fossil reefs. Facies 28, 109 – 114. Vogel, K., 1997. Bioerosion in rezenten Riffbereichen: Experimente vor Inseln Bahamas und des Grossen Barriereriffs. Natur und Museum 127, 198 – 209. Vogel, K., Keine, W., Gektidis, M., Radtke, G., 1996. Scientific results from investigations of microbial borers and bioerosion in reef environments. In: Reitner, J., Neuweiler, F., Gunkel, F. (Eds.), Global and Regional Controls on Biogenic Sedimentation I. Reef Evolution. Research Reports. Go¨ttinger Arbeiten aus dem Geologische – Paleontologischen, vol. Sb2, pp. 139 – 143. Vogel, K., Balog, S.-J., Bundschuh, M., Gektidis, M., Glaub, I., Krutschinna, J., Radtke, G., 1999. Bathymetric studies in fossil reefs, with microendoliths as paleoecological indicators. Profil 16, 181 – 191. Walter, L.M., Burton, E.A., 1990. Dissolution of recent platform

A.M. Smith, C.S. Nelson / Earth-Science Reviews 63 (2003) 1–31 carbonate sediments in marine pore fluids. American Journal of Science 290, 601 – 643. Walter, L.M., Morse, J.W., 1984. Reactive surface area of skeletal carbonates during dissolution: effect of grain size. Journal of Sedimentary Petrology 54, 1081 – 1090. Walter, L.M., Morse, J.W., 1985. The dissolution kinetics of shallow marine carbonates in seawater: a laboratory study. Geochimica et Cosmochimica Acta 49, 1503 – 1513. Warme, J.E., McHuron, E.J., 1978. Marine borers: trace fossils and their geological significance. In: Basan, P.B. (Ed.), Trace Fossil Concepts. SEPM Short Course No. 5, Tulsa, OK, pp. 67 – 118. Whittle, G.L., Kendall, C.G.St.C., Dill, R.F., Rouch, L., 1993. Carbonate cement fabrics displayed: a traverse across the margin of the Bahamas Platform near Lee Stocking Island in the Exuma Cays. Marine Geology 110, 213 – 243. Wollast, R., 1994. The relative importance of biomineralization and dissolution of CaCO3 in the global carbon cycle. In: Doumenge, F. (Ed.), Past and Present Biomineralization Processes, Considerations about the Carbonate Cycle. Bulletin de l’Institut Oce´anographique, Monaco, no. spe´cial 13, pp. 13 – 34. Young, H.R., Nelson, C.S., 1985. Biodegradation of temperatewater skeletal carbonates by boring sponges on the Scott shelf, British Columbia, Canada. Marine Geology 65, 33 – 45. Young, H.R., Nelson, C.S., 1988. Endolithic biodegradation of cool-water skeletal carbonates on Scott shelf, northwestern Vancouver Island, Canada. Sedimentary Geology 60, 251 – 267. Abigail M. Smith received her BA degree in Geology and Biology from Colby College, Waterville, ME, USA before going on to study at MIT and Woods Hole. Her MS in Earth, Atmospheric, and Planetary Sciences from MIT was the beginning of her research into temperate carbonates. In 1992, she completed her doctoral research at the University of Waikato, Hamilton, New Zealand, on bryozoans in temperate carbonate sediments. She is now Senior Lecturer in Marine Science at the University of Otago, Dunedin, New Zealand. Her research into temperate marine shelf carbonates includes studies of taphonomy, production, sedimentation, and geochemistry of these cool-water carbonates, particularly of bryozoans. She has been on the council of the New Zealand Marine Sciences Society and is currently treasurer of the International Bryozoology Association.

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Cam Nelson received BSc and BSc(Honours) degrees in Geology at Victoria University, Wellington, New Zealand. He then lectured in the Department of Geology at the University of Auckland where he received his PhD before joining the Department of Earth Sciences at the University of Waikato in Hamilton, New Zealand, in 1971 as its founding geological staff member. He was department Chairperson from 1988 to 1996, and has been a Professor since 1991. His research interests are in sedimentary and marine geology, including non-tropical carbonate facies, and Cenozoic paleoceanography and paleoclimatology of the southwest Pacific region. He is past President and office holder of the Geological Society of New Zealand, and was elected a Fellow of the Royal Society of New Zealand in 1994.