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The Geological Society of America Special Paper 436 2007
Carbonate-platform facies in volcanic-arc settings: Characteristics and controls on deposition and stratigraphic development Steven L. Dorobek* Department of Geology & Geophysics, Texas A&M University, College Station, Texas 77843, USA ABSTRACT Shallow-marine carbonate facies from volcanic-arc settings provide an important, but commonly overlooked, record of relative sea-level change, differential subsidenceuplift, paleoclimate trends, and other environmental changes. Carbonate strata are thin where volcanic eruptions are frequent and voluminous, unless shallow, bathymetric highs persist for long periods of time and volcaniclastic sediment and erupted materials are trapped in adjacent depocenters. Carbonate platforms and reefs can attain significant thickness, however, if subsidence continues after volcanic activity ceases or the volcanic front migrates. The areal extent of shallow-marine carbonate sedimentation is likewise affected by differential tectonic subsidence, although carbonate platforms are most laterally extensive during transgressive to highstand conditions and when arc depocenters are filled with sediment. Tectonic controls on shallow-marine carbonate sedimentation in arc depocenters include (1) coseismic fault displacements and associated surface deformation; (2) longwavelength tectonic subsidence related to dynamic mantle flow, flexure, lithospheric thinning, and thermal subsidence; and (3) large-scale plate deformation related to local conditions of subduction. Depositional controls on carbonate sedimentation in arc depocenters include (1) the frequency, volume, and style of volcanic eruptions; (2) accumulation rates for siliciclastic-volcaniclastic sediment; (3) the frequency, volume, and dispersal paths of erupted material; (4) (paleo)wind direction, which influences both carbonate facies development directly and indirectly by controlling the dispersal of volcanic ash and other pyroclastic sediment, which can bury carbonate-producing organisms; (5) the frequency and intensity of tsunami events; and (6) volcanically or seismically triggered mass-wasting events, which can erode or bury carbonate strata. Regarding platform morphologies in arc-related settings, (1) fringing reefs or barrier reef systems with lagoons may develop around volcanic edifices throughout the long-term evolution of volcanic arcs; (2) local reefs and mounds may build on intrabasinal, fault-bounded highs within underfilled forearc, intra-arc, and backarc basins; (3) isolated platforms with variable platform margin-to-basin transitions are
*Current address: Maersk Oil & Gas, Esplanaden 50, DK 1263 Copenhagen, Denmark. Dorobek, S.L., 2007, Carbonate-platform facies in volcanic-arc settings: Characteristics and controls on deposition and stratigraphic development, in Draut, A., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. XXX–XXX, doi: 10.1130/2007.2436(04). For permission to copy, contact
[email protected]. ©2007 The Geological Society of America. All rights reserved.
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Dorobek common in “underfilled” and tectonically active depocenters; and (4) broad ramps and rimmed carbonate shelves are typically found in tectonically mature and sediment-filled depocenters. Keywords: carbonate platforms, reefs, carbonate facies, subsidence, basins, depocenters, carbonate stratigraphy.
INTRODUCTION Shallow-marine carbonate facies may constitute an important part of the stratigraphy that forms in volcanic-arc settings, although they have received relatively minor attention from geoscientists (cf. Dickinson, 1995, 2001; Smith and Landis, 1995; Soja, 1996; Wilson and Boxence, 1997; Nunn, 1998a; Wilson, 2002; Wilson and Lokier, 2002; Wilson and Vecsei, 2005; Bosence, 2005). Shallow-marine carbonate facies in volcanic-arc settings are important recorders of sea-level changes, differential subsidence across arc depocenters, paleoclimate, and other environmental conditions. Thus, careful analysis of these strata is critical for reconstructing the complicated geological histories of volcanic-arc depocenters. Significant petroleum accumulations are also stored in carbonate reservoirs in some volcanic-arc depocenters (e.g., Indonesia backarc basin; Sharaf et al., 2005), although these rather limited examples may be associated with special petroleum-system conditions. This paper presents a review of the tectonic and depositional controls that influence shallow-marine carbonate systems associated with volcanic arcs along convergent margins. Examples of carbonate systems from modern and ancient settings are used to illustrate key principles, although it is important to note that relatively few, thoroughly documented examples exist from which to draw conclusions or construct predictive models. Detailed stratigraphic information for carbonate successions that formed in a variety of arc depocenters during different times in Earth history is simply not available. Instead, the main objective of this paper is to describe how tectonic deformation, other depositional controls, and siliciclastic and volcaniclastic sedimentation can influence shallow-marine carbonate facies development in arc settings. More complete documentation of carbonate successions from modern arc settings and from the rock record will undoubtedly lead to refinement of the basic models presented here, but I hope this paper serves as a general guide to the analysis of arcrelated carbonate systems. TECTONIC CONTROLS ON ARC-RELATED CARBONATE SYSTEMS Tectonic deformation along convergent plate margins strongly influences shallow-marine carbonate depositional systems in volcanic-arc depocenters. Differential tectonic subsidence or uplift in any tectonic setting can be attributed to (1) local, fault-controlled deformation, which can involve fault systems
that only cut upper crustal levels to truly lithosphere-scale fault systems; and (2) long-wavelength, basin-scale differential subsidence and uplift, which typically involve lithosphere-scale deformation processes (e.g., flexure, lithospheric thinning, or thermal subsidence) or dynamic mantle flow (i.e., dynamic topographysubsidence; cf. Gurnis, 1990a, b, 1991). Convergent margins with active subduction, however, are unique tectonic settings where coseismic displacements along great subduction faults may cause long-wavelength (>100 km) surface deformations that can significantly affect depositional systems on the upper (or overriding) plate. This style of differential tectonic subsidence and uplift is exclusive to active subduction zones and was dramatically illustrated during the December 2004 Sumatran earthquake (Fig. 1). These various scales of tectonic deformation typically interact simultaneously and in highly complex ways in volcanic-arc settings, making the subsidence and uplift histories of arc depocenters difficult to characterize, predict, and understand. Other tectonic characteristics of volcanic-arc settings can also affect carbonate sedimentation, such as large-scale aspects of subduction, plate motion, and plan-view geometries of plate margins. The following section first introduces large-scale tectonic controls on arc depositional systems and progresses to consider more local tectonic deformation and its effects on deposition in arc depocenters. Intra-Oceanic Island Arcs versus Continental-Margin Arc Systems Whether a volcanic arc forms as an intraoceanic or a continental-margin arc system is of first-order importance for shallow-marine carbonate systems. Intraoceanic island arcs form on the leading edge of an overriding oceanic plate. In contrast, continental-margin arc systems form where a slab of oceanic lithosphere is subducted beneath an upper plate of continental lithosphere. Arc-related basins along both types of plate margins may be similar, although the initial topography of the overriding plate and rheological differences between oceanic and continental lithosphere of the upper plate impose different conditions on basin development. Volcanism associated with intraoceanic arc systems initially occurs in submarine environments. Successive submarine eruptive events cause progressive shallowing of volcanic edifices until they eventually build to within the photic zone, and shallow-marine carbonate deposition can begin, which may take place after millions of years of eruptive activity. In
spe436-04 page 3 Carbonate-platform facies in volcanic-arc settings
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Figure 1. Distribution of coseismic uplift and subsidence across the Andaman Islands and Western Indonesian forearc after the December 2004 and March 2005 major earthquakes. Displacement along the subduction fault between the subducting Indian Ocean plate and the overriding Indonesian forearc propagated rapidly northward from the epicenters (indicated by stars) toward the Andaman Islands. Areas of coastal uplift are indicated by the stippled pattern. Areas of coastal subsidence are indicated by dark shading. Dashed line represents interpreted subsidence hinge line. Data obtained from SAR interferometry (InSAR) analyses. Modified from Japan Geographical Survey Institute Web site.
Convergent margins can be classified according to the regional styles of deformation that affect the upper plate (cf. Dewey, 1980). Various sectors of the upper plate can be characterized by extensional, contractional, strike-slip, or very minor (i.e., neutral) deformation. Extensional deformation can develop within (1) the thermally weakened basement of the arc massif, (2) the backarc region, or (3) the forearc region, especially where there is significant tectonic erosion during subduction (Clift et al., 1998). Most deforming volcanic arcs have at least local strikeslip fault zones, even where contractional or extensional styles of deformation are dominant. In compressional arcs, thrust and reverse faults typically form in the backarc region on the upper plate; these faults are typically antithetic to the dip of the subducting slab. Compressional retroarc foreland basin settings are not considered here, although carbonate facies in extensional backarc basins are discussed in a later section. It also is important to note that some dip-trending sections across subduction margins show compressional deformation in one segment of an arc system but other styles of deformation in other parts of the system. For example, the Sumatran forearc region is generally considered to be an “accretionary” forearc (Clift and Vannucchi, 2004), yet the arc itself is cut by the Sumatran Fault, which is a major, through-going, strike-slip fault system that extends along the entire length of Sumatra. Although the Sumatran forearc region is generally thought to be accreting, active zones of transtensional faulting cut at high angles across the forearc basin and partition the Sumatran forearc into multiple depocenters. These closely juxtaposed styles of tectonic deformation are described here simply to demonstrate that it can be inappropriate to make general characterizations about the overall tectonic style of a particular arc system without examining long-term deformation and subsidence histories across the entire system, from the trench to the backarc region, and understanding the kinematic history of the deformation. Accurate, large-scale kinematic analyses may be impossible in strongly tectonized, ancient arc settings. Margin Geometries, Relative Plate Motion, and Changes in Subduction
contrast, continental-margin arc systems may actually begin as subaerial settings and require either a significant eustatic rise or tectonic subsidence to submerge them to photic depths, when shallow-marine carbonate deposition can begin. Although both oceanic and continental lithosphere will bend in response to the growth of surface loads like volcanic arc systems, they have different flexural rigidities so their flexural response will differ. Continental lithosphere also may have more inherent strength variations than oceanic lithosphere because of its typically long and complicated geologic history. These variations in strength and the presence of preexisting faults can influence basin development in continental-margin arc systems.
Plate-motion vectors and the plan-view geometry of continental margins (i.e., before and during convergence and collision) also influence subduction or arc collision along specific margin segments (Dewey, 1980; Thomas, 1983; Bradley, 1989). The initiation of subduction or collision usually will be diachronous along strike where the motion of one plate is highly oblique to another plate margin. If either plate is rotating about a vertical pole, then the rate of subduction or shortening caused by collision typically increases with distance from the pole of rotation. The presence of salients or recesses along either plate margin will also influence which segments along either margin are first to collide.
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The angle of convergence or subduction also strongly influences the styles and distribution of deformation within the overriding plate. Where the relative motion of both plates is nearly perpendicular to the trench axis, and trench rollback is slower than the velocity at which the overriding plate approaches the subduction zone, compressional styles of deformation should dominate in the overriding plate (cf. Dewey, 1980). In contrast, oblique subduction characteristically produces some type of transtensional, transpressional, or pure strike-slip deformation in the overriding plate. The style and geographic location of upperplate deformation (e.g., within the arc-trench gap, the volcanic arc, or in the backarc region) and volcanic activity depend on (1) the angle of convergence; (2) the dip of the subducted slab; (3) whether any seamounts, active spreading-ridge segments, fracture zones, or other rough surface features on the subducted slab arrive at the subduction zone; (4) the degree of mechanical coupling between the subducting and overriding plates; and (5) the location and orientation of lateral strength variations or preexisting fault zones in the overriding plate, which might be reactivated during convergence. Other mechanical or rheological attributes of plates or microplates involved along the convergent margin may influence deformation in the overriding plate but probably in less significant ways. The timing of collision and styles of deformation associated with convergent margins will change if plate-motion vectors vary over time. Changes in the relative motion of plates also cause intraplate stress trajectories to be reoriented over time, and consequently the styles and patterns of deformation within plate interiors should evolve, especially where preexisting basement fault systems are reactivated by in-plane stress. Magmatic activity in the overriding plate will also be affected by changes in the nature of subduction (cf. Otsuki, 1989). Where the subducted slab has a shallow dip, the location of the volcanic front generally is farther inboard from the trench, or the margin may lack substantial volcanic and magmatic features altogether where “flat-slab” subduction occurs. The overriding plate may also be deformed where the dip of the subducted plate becomes very shallow (nearly horizontal) and there is mechanical coupling between the subducted plate and the underside of the overriding plate (e.g., Neogene deformation of the Chilean-Argentinean foreland; Jordan et al., 1983; Jordan and Allmendinger, 1986). If the dip of the subducted slab steepens over time, trench rollback and extension may occur across much of the arc (Uyeda and Kanamori, 1979; Otsuki, 1989). The plan-view geometry of the overriding plate may also control curvature of the subducting plate, downdip from the subduction zone, which in turn may influence deformation across backarc regions in the upper plate (cf. Cahill and Isacks, 1992). Where arc-continent or continent-continent collision has occurred, subduction will usually become very limited or may cease altogether. In these cases, the subduction zone may jump farther outboard and also change dip (or polarity) so that continued plate convergence can be accommodated by continued subduction of oceanic lithosphere. Where continental plates or arc
microplates meet along a convergent margin, some blocks might be “extruded,” or pushed laterally away from the zone of convergence (e.g., Tapponnier et al., 1982; Coward, 1990; Maynard et al., 1997), although the amount of extrusion and lateral translation of blocks away from the collision zone varies. Significant shortening may continue to occur along the convergence zone, but lateral translation of fault-bounded blocks may be required to accommodate additional convergence once the crust within the collision zone cannot be tectonically thickened anymore. Regional strike-slip-fault networks that bound the translated blocks accommodate the extrusion. The sense of slip on these faults may change over time as one block progressively plows into the adjoining plate and more inboard blocks move laterally to accommodate additional shortening in the collision zone. Tectonic Drivers of Long-Wavelength Differential Subsidence in Arc Settings Volcanic-arc settings are characterized by rapidly changing and complexly distributed subsidence patterns, which reflect the interactions of multiple tectonic drivers that operate along subduction margins. These drivers of differential long-wavelength subsidence generally involve deformation or long-term rheological changes of the entire lithosphere, and include flexure, thermal subsidence, magmatic underplating, and thinning in extensional arcs (Fig. 2). Dynamic mantle flow can also influence long-wavelength subsidence patterns in the upper plate (i.e., dynamic topography; Gurnis, 1990a, b, 1991). Long-wavelength subsidence patterns will influence the distribution, plan-view dimensions, and ultimate thickness of carbonate platforms, with the most widespread development of carbonate facies generally occurring during waning stages of subduction, when tectonic subsidence rates are progressively slowing and arc depocenters are more likely to become filled with sediment. Deformation in the forearc region, especially within the subduction complex, can result in other types of long-wavelength subsidence or uplift that are not related to lithospheric deformation. The subduction complex can deform in multiple ways, including complex internal thrusting, backthrusting, normal faulting, and gravitational sliding. These different types of deformation are thought to occur as the subduction complex tends to maintain a wedge shape or “critical taper” (Davis et al., 1983; Dahlen et al., 1984). Underplating and subduction erosion also can cause long-wavelength surface deformation across the top of the subduction complex (Fig. 3). Deformation of the subduction complex can, in turn, cause differential subsidence in the outer part of the forearc basin because the subduction complex commonly forms the seaward limit of this depocenter. This differential subsidence affects shallow-marine carbonate deposition in the outer part of the forearc basin when the basin is filled with sediment and the basin floor is within the photic zone. Along convergent margins where the subduction complex is absent or poorly developed (e.g., with sediment-starved trenches or where subduction erosion dominates), depositional
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Carbonate-platform facies in volcanic-arc settings DS Fsc
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Figure 2. Causes of long-wavelength (>100 km) tectonic subsidence across the overriding plate for different types of convergent margins. Flexural subsidence patterns (indicated by dashed black lines): Fsc—flexural subsidence of subducted plate caused by load of the subduction complex; Ffbf—flexural subsidence of forearc region caused by sediment load within the forearc basin; Fbabf—flexural subsidence of backarc region caused by sediment load within the backarc basin; Frfb—flexural subsidence of retroarc foreland basin region caused by sediment load and bordering thrust belt. Thermal subsidence patterns (indicated by dashed lines): Tva—thermal subsidence caused by cooling magmatic arc, after arc becomes extinct; Tbaoc—thermal subsidence of oceanic crust created along backarc spreading center; Tra—thermal subsidence caused by cooling of remnant magmatic arc in extensional backarc settings; Tba—thermal subsidence caused by cooling of rifted backarc continental or arc crust. DS—long-wavelength dynamic subsidence across retroarc settings (indicated by solid black line). In-plane stress may also affect any flexural subsidence patterns, depending on whether stress is compressive or tensile (in-plane stress effects are not shown here). Note how various tectonic subsidence mechanisms may interfere with subsidence patterns caused by other drivers. Patterns are highly generalized. Any of these drivers of long-wavelength subsidence can be counter-affected by shortening or tectonic inversion across the overriding plate. See text for additional discussion. (A) Intra-oceanic volcanic arc with old oceanic crust trapped in backarc region. There may be little or no thermal subsidence of the old, trapped oceanic crust. (B) Backarc region underlain by newly created oceanic crust formed by backarc spreading axis. (C) Compressional backarc region (retroarc foreland basin system) underlain by continental crust. (D) Extensional backarc region underlain by rifted continental crust. Synrift subsidence due to crustal thinning is not shown; only postrift thermal subsidence is shown for the extensional retroarc basin.
gradients from the arc to trench tend to be steep for long periods of time, and shallow-water carbonate deposition will be limited mostly to fringing reefs and narrow platform systems around the volcanic arc. Fault-Related Surface Deformation in Arc Depocenters Displacement along individual faults in arc settings can cause shorter wavelength surface deformation. This surface deformation dramatically influences carbonate deposition if the deformation occurs in shallow-water environments. Although
a thorough description of surface deformations caused by fault displacements is beyond the scope of this paper, typical coseismic surface deformations are described in Table 1. Important aspects of these surface deformations include the following: 1. Coseismic surface deformations are geologically instantaneous and can amount to several meters where fault displacements are greatest, which is typically found at the midpoint of fault surfaces along which displacement occurs. 2. Recurrence intervals of earthquakes that cause significant surface deformation in arc depocenters may be several
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A
uplift of trench-slope break caused by back-rotation of imbricate thrust sheets; local uplift patterns controlled by individual thrust slivers
B
uplift of subduction complex over zone of underplating; uplift geometry controlled by geometry of duplex system
C
uplift of trench-slope break caused by back-thrusting as subduction complex attempts to maintain critical taper; local fault-related folding
decades to several thousand years. Even with comparatively long recurrence intervals, surface subsidence and uplift can accrue to many tens or even hundreds of meters within a few thousand or tens of thousands of years, which will have obvious effects on carbonate facies development. Longer earthquake recurrence intervals may allow carbonate platforms to “recover” and fill new accommodation space that was created during an earthquake. 3. Coseismic displacements on great subduction faults can cause large areas of carbonate platforms and reefs in arc depocenters to subside or to be instantly uplifted above sea level (Fig. 1). Coastal and intertidal environments will be most greatly affected. These events generally affect the forearc region only, where mechanical coupling between the subducting and overriding plates is strongest. 4. Coseismic fault displacements can generate tsunamis or cause mass wasting that can obliterate or bury large areas of shallow-marine carbonate deposition. 5. Some faults may creep aseismically, and although their related surface deformations may slowly develop, significant amounts of uplift or subsidence can still accrue over time. Aseismic fault creep may be most common in subduction complexes where faults cut weakly lithified, highly sheared, fluid-rich materials. Active faults in arc depocenters can control carbonate-platform morphology, platform margin-to-basin profiles, and lateral facies changes during deposition. Active faults can segment preexisting carbonate platforms or control the location of carbonate
backthrusting
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subsidence caused by tectonic erosion or sediment subduction; compaction has similar effect
tectonic erosion beneath subduction complex
sediment subduction
Figure 3. Possible mechanisms for internal deformation and evolution of subduction complexes, with associated short- to intermediatewavelength surface deformation. Local surface deformations commonly are expressed as complex patterns of uplift, which are superimposed on longer-wavelength patterns of tectonic subsidence. Figure modified from Dickinson (1995). (A) Accretionary offscraping of successive imbricate thrust panels (Seely et al., 1974). For growing subduction complexes, back-rotation of imbricate thrusts causes progressive uplift of crest of subduction complex (i.e., trench-slope break region). There also may be local fault-related folding, which creates antiformal surface deformations. (B) Selective subduction and underplating of thrust duplexes (Sample and Fisher, 1986). Underplating at the base of the subduction complex can cause uplift over the zone of underplating. Structural styles related to underplating are often poorly imaged on seismic profiles but probably are most likely expressed as a duplex zone. (C) Backthrusting of subduction complex toward forearc basin (Silver and Reed, 1988). Uplift of trench-slope break and distal part of forearc basin may occur as the subduction complex maintains a critical taper. Backthrusting in arcward side of subduction complex generally causes uplift in trench-slope break region, with local surface deformation caused by fault-related folding. (D) Denudational normal faulting. Both regional subsidence and local fault-controlled subsidence may occur across trench slope as the subduction complex maintains a critical taper via extensional deformation. (E) Subduction erosion along décollement beneath subduction complex (cf. von Huene et al., 1982). Subduction erosion or sediment subduction along underside of subduction complex creates subsidence across subduction complex. Compaction of subduction complex may have similar effects.
spe436-04 page 7 Carbonate-platform facies in volcanic-arc settings Table 1. TYPICAL COSEISMIC SURFACE DEFORMATIONS FOR DIFFERENT FAULT TYPES Fault type Fault length Surface upliftinvolved subsidence Normal faults 5–50 km 0–3 m Reverse/thrust faults 5–100 km 0–5 m Subduction faults 100–1000+ km 0–5 m Strike-slip faults: 1–50 km 0–3 m Transfer fault 50–500+ km 0–3 m Major intraplate or plate boundary
7
The bathymetry of arc depocenters can be rapidly modified by subaerial or submarine deposition of volcanic materials or large-scale mass-wasting events. Arc settings are also seismically active during active subduction, and large-magnitude earthquakes commonly generate significant ground displacements, so slope failure occurs frequently. Although erosional and depositional modifications to shallow-water depositional profiles are relatively common in arc depocenters, they are (1) episodic and thus difficult to predict and correlate in stratigraphic successions, (2) variable in lateral extent, and (3) involve highly variable amounts of rock materials.
and stratal patterns (Caron et al., 2004). Elongated, tide- and storminfluenced balanid (barnacle) shoals grew on local antiformal highs near the crest of the subduction complex. In contrast, arc-attached, low-gradient shelf and ramp profiles were built along the arc side of the forearc basin. Carbonate strata are more sheetlike along the flank of the arc, are generally transgressive in character, and were dominated by epifaunal bivalves. These characteristics are thought to reflect a higher flux of sediment from the adjacent arc. The transport mechanisms and volume of material involved in sediment gravity flows will control the size of carbonate platforms or reefs that might be buried during a single event. The frequency of catastrophic depositional events within a particular segment of the arc will determine whether adjacent carbonate platforms and reefs can recover from a series of events. If relatively low volumes of volcanogenic material are involved in any catastrophic event, and the events occur with low frequency or are widely dispersed along strike, then carbonate platforms and reefs may recover and even utilize lava flows or volcaniclastic facies as substrates. Mass wasting can also occur at any time along steep volcanic edifices, although these catastrophic events are triggered more often by volcanic earthquakes or by inflation of the volcano’s surface as magma ascends into subsurface magma chambers prior to an eruption. Antecedent drainage networks on the flanks of a volcano may control dispersal of low to moderate volumes of volcanogenic material, although dispersal of large-volume, density-stratified pyroclastic flows are generally less affected by antecedent topography. The flow paths of only the basal, highdensity part of pyroclastic flows may be diverted by antecedent topography, such that the resultant coarse-grained deposits are confined to ravinement floors on the flanks of a volcano. These high-density flows then spread laterally, where they become unconfined in lower hillslope regions where antecedent valleys are less incised and surface gradients flatten. In contrast, the lowdensity upper parts of pyroclastic flows are commonly stripped from the basal, high-density part of the flows. The low-density parts of pyroclastic flows can flow over topographic obstructions and lay down more areally extensive deposits. Low-density pyroclastic flows can also flow for several kilometers across seawater. Lava flow paths will also be controlled by antecedent topography, unless the volume of lava exceeds the ability of valleys and hillslope ravinements to contain the flows.
Sediment Flux to Arc Depocenters
Underfilled versus Filled Arc Depocenters
Sediment flux from the volcanic arc or uplifted arc-massif rocks can have many effects on carbonate production in adjacent depocenters. High sediment flux can completely overwhelm carbonate-producing benthic organisms. Lesser amounts of sediment flux might lead to mixed carbonate-siliciclastic-pyroclastic successions. The caliber of sediment supply from volcanic-arc and arc-massif rocks can also determine the types of carbonate sediment produced. For example, late Pliocene carbonate facies in the New Zealand forearc show major lateral changes in sediment types
Sedimentary basins are often described as “underfilled” or “filled” (cf. Covey, 1986; Flemings and Jordan, 1990; Jordan, 1995), which describes whether depositional systems deposit sediment into a depocenter or carry sediment past a former depocenter. Although a thorough discussion of this concept is beyond the scope of this paper, it is useful for understanding where and when widespread carbonate platforms might form in arc depocenters. Although unique conditions of subduction can determine subsidence and basin-filling patterns in arc systems, arc depocenters
facies tracts wherever displacement rates exceed the ability of carbonate facies to fill newly generated accommodation space on either side of a fault. Carbonate lithofacies deposited around faults that were active during deposition may contain greater amounts of gravity-flow deposits, tsunamites, or other catastrophic-event beds. DEPOSITIONAL CONTROLS ON CARBONATE SEDIMENTATION IN ARC SETTINGS The accumulation of noncarbonate sediments and eruptive products has important effects on carbonate deposition in arc depocenters. Various depositional and erosional processes that are characteristic of arc settings can also affect carbonate deposition. The relative influence of these factors on carbonate deposition may depend on the developmental stage of the arc system. Erosional and Depositional Modifications to Bathymetric Profiles
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are generally more likely to become filled (1) during waning stages of subduction when tectonic subsidence slows down, (2) in arc depocenters with high sediment accumulation rates, or (3) in arc depocenters that begin near base level at the onset of subduction. Previously deposited siliciclastic sediment and erupted materials in filled depocenters can provide broad substrates for shallow-marine carbonate deposition. Areally extensive carbonate platforms most commonly form when rising sea level floods these substrates. In addition, tectonic surface deformations become more important for shallow-water carbonate depositional systems in filled depocenters. If a filled depocenter is submerged to water depths that are within wave base and typical tidal ranges (2 m locally. A regional, arc-parallel hinge line apparently separates uplifted from subsided parts of the forearc, with uplifted regions on the trenchward side of the forearc basin. This uplift may be preserved in the rock record as a regional disconformity within shallow-water strata. It would be difficult to recognize whether these disconformable surfaces in forearc strata were related to coseismic surface deformation or high-frequency sea-level falls. Regardless, the recent Sumatran earthquake provides an important reminder of the rapid uplift and subsidence that can occur in forearc settings. Other modern forearc basins where Quaternary carbonate facies are found include (1) the forearc basin along the southeastern side of North Island, New Zealand (Kamp and Nelson, 1988), and (2) the island of Espiritu Santo, Vanuatu (Wells, 1988; Johnson and Greene, 1988; Greene et al., 1988; Cabioch et al., 1998). North Island, New Zealand. In the modern forearc basin of North Island, New Zealand, thin veneers of carbonate sediment are being deposited over only ~3% of the forearc region (Kamp and Nelson, 1988). This limited area of modern shallow-water carbonate sedimentation reflects the combined deleterious effects of the relatively cool-water conditions of offshore eastern New Zealand, the high-energy wave-swept conditions of this forearc setting, and the flux of siliciclastic sediment from North Island.
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Figure 8. (A) Schematic cross sections that depict the effect of relative sea-level changes on coral growth. The outermost solid-black growth band depicts the most recent coral growth band. Solid gray bands are dead coral growth bands that grew under the same sea-level conditions as the living coral. Dashed growth bands are dead bands that grew under different, prior sea-level conditions. (a) Hemispherical coral head below highest level of survival (HLS), with a new skeleton accreting in a concentric fashion on the outside of the head. (b) The same coral head under the same sea-level conditions after it has reached HLS. No additional upward growth is possible, although lateral accretion continues along the sides of the coral head that are below HLS. (c) “Hat” morphology of a coral head that grew up to HLS, but then underwent a drop in HLS. The part of the coral exposed above HLS has died. Lateral growth below the new HLS develops a lower outer rim around a higher center. The elevation difference between the two flats is a measure of the amount of emergence. (d) “Cup” morphology of a microatoll that underwent a sea-level rise after the coral had been growing at HLS. The coral grows upward toward the new HLS, constrained only by its growth rate. Upward and outward growth over the old HLS surface produces a raised outer rim, indicative of submergence. The elevation of the new HLS is not recorded by the coral until the coral grows up to it. From Zachariasen et al. (1999). (B) Cross section of a microatoll from the Indonesian forearc. The X-rayed thin slab reveals a clear record of annual growth bands that expanded radially outward (from left to right) at ~1 cm/yr. The HLS of the coral during the past 35 yr is recorded in the topography of the coral’s upper surface. The arrows track the apparent rise of sea level in the 1960s and its subsequent apparent fall. Growth patterns in corals such as these can provide high-resolution records of true sea-level change or the effects of incremental coseismic subsidence or uplift. From Sieh et al. (1999).
spe436-04 page 15 Carbonate-platform facies in volcanic-arc settings Beyond the modern shelf, carbonate sediments (30,000
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now form the basement rocks of the present-day Vanuatu forearc region. A subduction complex is absent along this margin owing to its young age and the limited amount of sediment delivered to the trench. The central part of the Western Belt, including Espiritu Santo, has been actively uplifting throughout Quaternary time owing to subduction of the d’Entrecasteaux Ridge at this location. The uplift history of Espiritu Santo is recorded by a series of emergent and notched Holocene reef terraces that were deposited on the uplifting Oligocene–Miocene volcanic substrate (Fig. 9; Cabioch et al., 1998). Although we know the present-day tectonic setting of Espiritu Santo, and the Holocene reef terraces should be described as having formed in a forearc setting, it might be difficult to make the same tectono-stratigraphic interpretation if similar carbonate terraces were found in the rock record. More likely, these reef terraces would be described as “intra-arc basin” deposits (see discussion below about intra-arc carbonate facies), simply because they are found stratigraphically above deformed, arc-related rocks. Seismic profiles from offshore Espiritu Santo also show drowned Quaternary carbonate reefs and reef caps (1– 6 km across, some reefs up to 300 m relief, reef tops at ~300 m water depth) and “mounded” sediment bodies (375–450 m water depth) (Fig. 10; Johnson and Greene, 1988; Greene et al., 1988), which formed on submerged and complexly deformed parts of the relict arc basement. The locations of some reefs and shelf margins are probably fault controlled. These highly variable patterns of Quaternary carbonate sedimentation around Espiritu Santo reflect the complex differential subsidence and uplift that can occur across young, incipiently formed forearc regions. In this case, Espiritu Santo represents subaerial remnants of an old volcanic arc that were deformed and became basement for the forearc region of a younger volcanic arc. It is worth noting that Quaternary carbonate facies apparently are extremely rare within forearc basins along the westfacing, convergent margins of North and South America, even
6 ka reef flat
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Figure 9. Cross section of the uplifted Tasmaloum fringing reef, Vanuatu, showing the 6, 8, 10, 12, 14, 16, 18, and 20 ka time lines. Only the most characteristic dates are reported (those in italics are 14C dates converted to calendar years). A 6 ka reef flat is now 40 m above sea level, which shows the rapid uplift rates that have characterized this forearc region over the last 6000 yr. From Cabioch et al. (1998).
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spe436-04 page 17 Carbonate-platform facies in volcanic-arc settings
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Figure 10. (A) Main geologic and tectonic features of the New Hebrides Arc. SMFZ—Santa Maria fracture zone; ABFZ—Aoba fracture zone; AMFZ—Ambrym fracture zone; EFZ—Epi fracture zone; VFZ—Vulcan fracture zone. Modified from Greene et al. (1988). (B) Main geologic and tectonic features of the central New Hebrides Arc, Espiritu Santo region. Basement rocks of Espiritu Santo consist of late Oligocene to middle Miocene volcanic and volcaniclastic rocks erupted from a paleo-volcanic arc (Western Belt rocks) that formed above a now-extinct, westward-dipping subduction zone. During late Miocene time, however, the subduction zone switched to the western side of the New Hebrides Arc (present-day New Hebrides Trench), and an east-dipping subduction zone developed beneath arc rocks of the Western Belt, which includes Espiritu Santo. The active volcanic arc to this subduction zone consists of volcanic centers that make up the Central Chain, to the east of the Western Belt. Thus, Western Belt rocks now constitute the basement of the modern forearc region of the Central Chain, and Espiritu Santo is in a transitional tectonic state. Depocenters that formed during extension of the older Oligocene–Miocene arc (Western Belt rocks) might be considered intra-arc basins. The volcanic centers of the Western Belt are no longer active, however, so the present-day depocenters should be considered part of the forearc of the Miocene–Holocene volcanic arc (i.e., the Central Chain). Modified from Greene et al. (1988). (C) Interpreted seismic-reflection profile across the East Santo and North Aoba basins, New Hebrides arc system. Note Quaternary pinnacle reefs–buildups along the outer, fault-controlled margin of the East Santo shelf. Modified from Greene et al. (1988).
those margin segments that lie within tropical climatic zones. The limited amounts of modern carbonate sedimentation along these margins probably reflects the strong coastal upwelling systems that bring cold, nutrient-rich bottom water onto the shallow, narrow shelf areas that border the western sides of these forearc basins. In contrast to modern continental margins, extensive Paleozoic carbonate platforms formed along the eastern side of the paleo–Pacific Ocean because marginal oceanographic barriers (e.g., island arcs, orogenic belts, and microplates) may have protected the carbonate platforms from the deleterious effects of upwelling (Whalen, 1995). Quaternary carbonate reefs and shallow-platform facies are also found in some relict forearc basin settings where subduction is no longer active but where the basin’s bathymetry and structural features were inherited from previous times of active subduction. In some ways, these basins might be considered forearc successor basins. For example, the Indispensable Basin of the eastern Solomon Islands has been interpreted as either a relict forearc basin or the leading edge of the partially subducted Ontong-Java Plateau (Vedder and Bruns, 1989). The Ontong-Java Plateau resisted southwestward-directed subduction when it collided with the Solomon Islands arc in middle to late Miocene time, forcing a change in subduction polarity and initiation of a younger, northeastward-directed subduction zone beneath the southwestern side of the Solomon Islands (Vedder and Bruns, 1989). Tectonic structures and remnant bathymetry along the relict forearc side of the Solomon Islands continue to influence modern sedimentation there, even ~15 m.y. after the change in subduction polarity. For example, seismic profiles along the southwestern slope of the Indispensable Basin show several drowned Pleistocene(?) reefs at ~300–500 m water depths along a steep, faulted(?) escarpment just southeast of Santa Isabel Island (Niem, 1989). The steep bathymetric gradient across the forearc region and the low relief and relatively small surface area of the relict volcanic arc suggest that this relict forearc basin will probably remain underfilled for a long time into the future, thus limiting the potential surface area for shallow-marine carbonate sedimentation. Bathymetric profiles around offshore parts of Viti Levu, Fiji, are structurally controlled and record older Neogene to possibly Quaternary deformation. Faults that cut the older Neogene forearc basin and
frontal arc massif influenced the location, dimensions, and morphology of Neogene to Holocene reefs across the area (Fig. 11; Rodd, 1993). Ancient examples. Ancient examples of forearc-basin carbonate sequences include (1) Paleogene carbonate platforms on transpressional ridges that transected the forearc basin of Costa Rica (Krawinkel and Krawinkel, 1996); (2) Oligocene to Miocene reefs and platform facies (Tau, Wailotua, and Qalimare Limestones) in the Outer Melanesian forearc basin on the island of Viti Levu, Fiji (Hathway, 1994, 1995; Rodd, 1993); (3) Miocene Rethymnon Formation, eastern Crete (Pomoni-Papaioannou et al., 2003); (4) Miocene mounded carbonate buildups to carbonate shelf systems in several sub-basins of the Indonesian forearc (Beaudry and Moore, 1985; Matson and Moore, 1992; van der Werff, 1996); and (5) Pliocene “cool-water” skeletal limestones, Hawke’s Bay, New Zealand (Kamp et al., 1988; Ballance, 1993; Caron et al., 2004). The present-day Terraba belt represents the inverted Tertiary forearc region of Costa Rica (Krawinkel and Krawinkel, 1996). This basin underwent three stages of basin evolution: (1) an initial transpressional phase during Late Cretaceous–Eocene time, which created the transpressional ridges that served as substrates for carbonate-platform facies; (2) a transtensional phase during Oligocene to early Miocene time, when the forearc region was affected by transtensional fault systems and was subsiding rapidly, causing the carbonate platforms to drown; and (3) a second transpressional phase of deformation from middle Miocene to Holocene time. Thus, tectonic deformation across the Costa Rican forearc basin initially created the substrates for carbonate sedimentation but ultimately caused the platforms to drown. Upper Oligocene to Miocene carbonate intervals exposed on the island of Viti Levu, Fiji, and imaged on seismic profiles from adjacent offshore areas consist of thin reefs or platform sequences that are interstratified with volcaniclastic strata (Figs. 11, 12; Hathway, 1994, 1995; Rodd, 1993). The Qalimare Limestone, which crops out on Viti Levu, is at least 300 m thick and consists of coral-algal reefs and mounded facies that formed on the edge of a shallow platform in front of the arc massif. Seismic profiles across the Bligh Water Basin from the northern offshore shelf of Viti Levu have intrabasinal highs that served as substrates for
spe436-04 page 18
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spe436-04 page 19 Carbonate-platform facies in volcanic-arc settings
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Figure 11. (A) Geologic map of Fiji. Most of the platform areas are shallow-water carbonate environments, although some platform areas are presently at subphotic depths (i.e., drowned or incipiently drowned). Also note the steep, typically fault-bounded edges of most platform areas, as indicated by solid black lines. The islands that make up Fiji represent the eroded and faulted remnants of an older volcanic arc–forearc basin system. Bathymetry in meters. From Rodd (1993). (B) Geologic cross section across Fiji and adjacent platform areas. This section shows various isolated platforms and buildups in the relict Bligh Water forearc basin. Note how faults controlled platform locations, dimensions, and morphology. From Rodd (1993).
local late Oligocene to middle Miocene reef development within the basin. The Miocene Rethymnon Formation (300 m above sea level) along the eastern flank of the basin; (2) successively younger Pliocene carbonate strata, shifted laterally toward the central axis of the forearc basin, especially along the more uplifted eastern flank of the basin; (3) older Pliocene carbonate sand bodies concentrated along the eastern and western flanks of the forearc basin, grading laterally into siliciclastic facies in the basin axis, whereas younger carbonate intervals are thinner and more sheetlike in character; and (4) unconformities within the Pliocene section that merge, becoming composite unconformities on the flanks of the forearc basin (Kamp and Nelson, 1987, 1988; Caron et al., 2004). General tectono-stratigraphic model for carbonate facies in forearc regions. Although detailed studies of carbonate strata deposited across forearc regions are relatively rare, previous studies of the tectono-geomorphologic evolution of forearc regions provide insight into the possible stratigraphic evolution of carbonate systems that might form there. The evolution of carbonate platforms and reefs across the forearc region will be influenced by (Fig. 6):
Sea level
Isolated platform on faulted high
arc uplift = increased sediment flux to forearc basin
c Ar
s ma
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Figure 12. Schematic tectonic and depositional model for southwestern Viti Levu during early to middle Miocene time. View approximately toward present northeast. Note how fault-controlled isolated platforms of the lower to middle Miocene Qalimare Limestone were deposited on the frontal side of the volcanic arc, which was rising isostatically because of crustal thickening. This in turn caused more erosion of the arc and greater flux of epiclastic sediment to the forearc basin. From Hathway (1994).
spe436-04 page 20 20
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Figure 13. Pliocene stratigraphy across the forearc basin of North Island, New Zealand. Measured stratigraphic sections at specific locales across the forearc basin are shown in the upper part of this figure; stratigraphic thickness in meters. Note the association of limestone with siliciclastic units and the progressive displacement of successively younger limestone units toward the basin axis. The cross section does not show all the structures distributed across this forearc basin and implies only broad, synformal folding over time. Most Pliocene limestone was deposited along the outer, trenchward side of the basin, although younger carbonate units shifted progressively westward over time, toward the arc. Also note the progressive decrease in dip of successively younger units (see lower cross-sectional view). These growth stratal patterns reflect local surface deformation above fault-propagation folds and more regional differential subsidence/uplift across the North Island forearc basin, which is expressed as the progressive westward shift of shallow-marine carbonate sedimentation across the basin. New Zealand stage names and their relative ages include: Wn—Nukumaruan (late PliocenePleistocene); Wm—Mangapanian (late Pliocene); Wp—Waipipian (middle Pliocene); Wo—Opoitian (early Pliocene); Tt—Tertiary. No vertical exaggeration. From Kamp and Nelson (1988).
1. Whether the subduction complex is accretionary or nonaccretionary. Accretionary subduction complexes grow laterally as trench sediments are added to the frontal part of the deforming wedge. Thus, accretionary subduction complexes commonly widen and shallow over time and provide ever-increasing potential surface areas for shallow-water carbonate sedimentation. 2. Differential uplift or subsidence across the forearc region. For example, if a formerly submerged subduction complex is uplifted and becomes emergent, any preexisting carbonate platforms or reefs across the crest of the subduction complex will become subaerially exposed, and carbonate depositional environments will be forced to
shift laterally off the flanks of the subduction complex. The total area available for shallow-water carbonate sedimentation may actually decrease, especially if uplift creates steep bathymetric gradients across the top of the subduction complex. If the upper surface of the subduction complex is below photic depths at the beginning, however, uplift may bring the surface into shallow-water depths and allow a carbonate factory to develop. If the top of the subduction complex is already emergent or at shallow water depths, continued uplift of the subduction complex may produce progradational to downstepping stratal patterns within carbonate platforms or fringing reefs that cap the subduction complex. In contrast,
spe436-04 page 21 Carbonate-platform facies in volcanic-arc settings regional subsidence of the subduction complex (owing to subduction erosion, flexure of the upper plate, or other tectonic causes) might cause preexisting carbonate platforms and reefs either to drown or to backstep toward any remaining shallow-water parts of the subduction complex. Similar complex chronostratigraphic and facies relationships may develop above other types of actively growing intrabasinal structural highs within the forearc basin or along the proximal arcward side of the forearc basin. 3. Transcurrent faults that can partition differential subsidence or uplift across the forearc region. Transcurrent faults that cut the forearc region are most common where oblique convergence occurs (e.g., Western Indonesian forearc). Transcurrent faults can partition parts of the forearc into antiformal highs or fault-bounded horst and graben segments, which control the location of platforms or isolated reefs (cf. Krawinkel and Krawinkel, 1996). Given the complex subsidence and uplift histories of forearc regions, carbonate-platform and reef strata that form across the forearc region during active subduction will likely be thin (5 km thick and consist largely of volcanic and volcaniclastic clasts, with lesser amounts of sandstone, mudstone, and shallow-marine carbonate facies. Platform- and reef-derived clasts (with dimensions up to several hundred meters) within these sedimentary breccias may provide the only evidence for former shallow-marine carbonate systems in many intra-arc basins (cf. Soja, 1990). Intra-Arc Basin Classification and Tectonic Models Intra-arc basins are depocenters within volcanic arcs and are underlain by arc-massif crust. Smith and Landis (1995) classified intra-arc basins as volcano-bounded, fault-bounded, or hybrid types. Volcano-bounded intra-arc basins are depocenters found between constructional volcanic edifices. Volcano-bounded basins associated with continental-margin subduction systems typically have only thin sediment fills, which Smith and Landis (1995) suggested are due to the generally high-standing character and low subsidence rates of their continental-crust substrates. In contrast, volcano-bounded basins from intraoceanic settings may accumulate thick volcanogenic strata, possibly because they subside more rapidly or begin as deep-water depocenters on oceanic crust. Erosion rates, however, are generally much higher in continental-margin arc systems than in intraoceanic arc systems, because continental-margin arc systems are largely subaerial features, whereas volcanic highs and intra-arc basins may be largely submerged in intraoceanic arc systems. Thus, erosion of volcanic highs also affects the amount of sediment delivered to intra-arc basins and the rate at which they are filled. Fault-bounded intra-arc basins are defined by fault networks. Fault patterns within most volcanic arcs and first-motion studies of intra-arc earthquakes indicate that extension or strike-slip deformation causes most fault-bounded intra-arc basins to form. Fault-bounded intra-arc basins typically have bounding faults that trend parallel to the volcanic arc, although some basins and their main bounding fault systems trend transverse to the volcanic arc. Fault-bounded intra-arc basins also may subside more rapidly than volcano-bounded depocenters (Smith and Landis, 1995). Transverse oblique-slip fault systems that delineate some intra-arc basins may accommodate block rotations along the arc axis, such as in the Aleutian forearc region (Geist et al., 1988). The curved trend of the Aleutian arc system, the oblique subduction that occurs along this convergent margin, and the strong coupling between the subducting and overriding plates may explain the block rotations. Differential subsidence of narrow crustal blocks between transverse faults may create submarine channels that funnel sediment gravity flows to the forearc basin or trench slope. Arc-parallel strike-slip faults can also accommodate significant lateral displacement of arc segments on either side of the
spe436-04 page 23 Carbonate-platform facies in volcanic-arc settings fault zone (e.g., Semangko or Sumatra Fault, Beck, 1983; Philippine Fault, Karig et al., 1986; Sarewitz and Lewis, 1991). Hybrid intra-arc basins have elements of both volcano- and fault-bounded basins. Hybrid basins are defined by Smith and Landis (1995) as depocenters with fault-defined margins, but where most of the topographic or bathymetric relief is defined by constructional volcanic features. Arc massifs and associated intra-arc basins typically undergo complex deformation histories and differential subsidence as a consequence of changes in subduction parameters (dip angle, obliquity, and rate of subduction), dynamic mantle flow, changes in surface loads, and magmatic activity. These large-scale tectonic and magmatic processes and resultant patterns of uplift and subsidence within the volcanic arc change over time scales on the order of 5–20 m.y. (Smith and Landis, 1995). As a result of the constantly changing tectonic and magmatic conditions that characterize the long-term history of most volcanic arcs, intra-arc basins may transition rapidly between volcano-bounded, fault-bounded, and hybrid types. All styles of deformation have been recognized within volcanic arcs and their related intra-arc basins. Extensional and strikeslip structures are generally more common within the arc massif than contractional structures, unless arc-continent collision has occurred. The history of deformation within intra-arc settings is poorly preserved because it is obscured and overprinted by multiple eruptive and deformation events, thick sediment accumulations may deeply bury and obscure fault networks, or the arc and its related basins are eventually incorporated into collisional orogenic belts and original tectono-stratigraphic relationships are destroyed. Many arc systems also record flips in subduction polarity such that former forearc regions may become backarc or interarc settings, or vice versa. Preexisting faults within the arc massif also may be reactivated by a different sense of displacement during subduction-polarity flips or other changes in subduction characteristics. These complicated tectonic histories may make it difficult to characterize the geometry, distribution, and kinematics of fault networks that develop within a volcanic arc. Extensional deformation is common where trench rollback causes the volcanic arc to split apart. An active seafloor-spreading ridge may form if the arc is completely stretched apart, leaving a remnant arc on the trailing side of the spreading axis and an active arc on the other side. In these cases, intra-arc rift basins evolve into an extensional backarc setting. Intra-arc rifting may occur faster than in intracontinental rift settings, with correspondingly much faster rates of synrift subsidence (Yamaji, 1990). Some volcanic arcs, such as the Marianas arc system, have been repeatedly split by intra-arc rifting and seafloor spreading (Hamilton, 1979). Other arcs have been split apart during a flip in subduction polarity; changes in subduction polarity may not occur synchronously along all segments of the convergent margin. These factors may create problems with basin classification and understanding of the tectonic processes that cause complex spatial and temporal patterns of uplift and subsidence. Examples of Cenozoic volcanic arcs that have been rifted apart during changes in subduction polarity include the Solomon Islands (Johnson, 1979; Kroenke,
23
1984; Bruns et al., 1989) and Vanuatu archipelagos (Carney et al., 1985; Greene et al., 1988). The fault-bounded basement highs that delineate intra-arc basins are parallel or oblique to arc segments. These structural highs typically serve as (1) minor sediment source areas if they are subaerially exposed, (2) substrates for shallow-marine sedimentation if they are only slightly submerged, or (3) bathymetric obstructions to sediment gravity flows if the structural highs are more deeply submerged. Although seismic data are limited from recently rifted arcs like the Solomon Islands and Vanuatu, the arc-parallel or arc-crossing highs that separate intra-arc depocenters within these arcs probably represent oblique transfer zones. Oblique dip-slip displacement on the faults that bound the intraarc highs allows extensional strain to be transferred between normal fault systems within the arc. These intra-arc fault-bounded highs may influence sedimentation within the arc for a long time after they form, especially if extension stops before the arc is completely rifted apart. Transtensional, transpressional, and pure strike-slip styles of deformation within the arc and intra-arc basins are more likely to develop where there is oblique subduction. As the convergence direction becomes more oblique or subparallel to the trend of the subduction system, pure strike-slip fault networks should become more common. Along some modern arc systems with highly oblique subduction angles, regional strike-slip faults may cut through and translate long segments of the volcanic arc (e.g., Sumatran or Semagang Fault along the island of Sumatra, western Indonesia; Philippine Fault, Philippine archipelago). Regional shortening within most volcanic arcs typically does not develop until arc-continent or continent-continent collision occurs and magmatic activity within the arc has long ceased. The limited amount of shortening within many volcanic arcs (and backarc regions) may indicate that mechanical coupling between the overriding and subducting plates diminishes with increasing distance from the trench (cf. Uyeda and Kanamori, 1979). Platform Types Most carbonate sequences within intra-arc basins consist of either fringing reefs constructed on volcanic or plutonic basement or barrier reefs with relatively narrow (100 m thick) before a catastrophic eruption or mass-wasting event buries the carbonate system. Mass-wasting events also might incorporate carbonate strata, causing large-scale disruption or transport of shallow-marine carbonate facies into deeper-water settings. Displaced blocks of reef and shallow-water carbonate facies within deep-water shale and volcaniclastic deposits may provide the only evidence of the former intra-arc platforms. There may be differences between intra-arc carbonate successions from intraoceanic and continental-margin arc systems. In intraoceanic arcs, shallow-marine carbonate facies will likely not appear within intra-arc basins until the volcanic arc has built itself into the photic zone. Volcano-bounded and hybrid intra-arc basins may dominate over fault-bounded basins in young intraoceanic volcanic arcs. The basal part of the basin-filling sequence in these basins will likely consist of deep-water volcaniclastic and pelagic-hemipelagic deposits. After volcanic activity has
Carbonate sedimentation also occurs in backarc settings on a variety of tectonic structures and bathymetric highs constructed by depositional processes. As in other arc depocenters, the areal extent, thickness, morphology, and internal stratigraphy of carbonate platforms and reefs will be determined by the tectonic history of the backarc and whether it is filled with sediment.
Carbonate Platforms and Reefs in Backarc Basins
Backarc-Basin Types and Tectonic Models Backarc basins form on the back side of subduction-related volcanic arcs (Fig. 16). Extensive geophysical surveys of many modern backarc regions over the last 20 yr or so have led to the recognition of three main types of backarc basins (Dickinson, 1974; Karig, 1983; Ingersoll, 1988; Marsaglia, 1995): (1) extensional backarc basins that form by rifting and seafloor spreading within or behind continental-margin or intraoceanic arcs, (2) remnant-ocean backarc basins that form by entrapment of old oceanic crust and are associated with intraoceanic subduction zones, and (3) compressional backarc basins, which are more commonly classified as retro-arc foreland basins (see discussion below). Marine backarc basins may be underlain by continental to transitional crust if they form by extension behind a continental-margin arc system. In contrast, backarc basins will be underlain by highly subsided oceanic crust if the basin forms along an intraoceanic subduction zone, and remnant oceanic crust is trapped behind the newly formed volcanic arc. Backarc basins also can be underlain by young oceanic crust if they form where rifting in the backarc region ultimately leads to seafloor spreading.
spe436-04 page 27 Carbonate-platform facies in volcanic-arc settings VOLCANO-BOUNDED INTRA-ARC BASINS platform types: • isolated platforms & buildups, commonly steep-sided • thin fringing platforms & reefs during early stages; complexly interstratified with volcanics, volcaniclastics • may be preferential development on windward side of arc • complex chronostratigraphic relationships between individual carbonate sequences • may be drowned platforms adjacent to actively growing platforms • platforms become larger & thicker as depocenters fill, but eruptions can have catastrophic effects on platforms substrates: • volcanic edifices & volcaniclastic fill
Volcano-bounded intra-arc basins
27
FAULT-BOUNDED INTRA-ARC BASINS platform types: • isolated platforms & buildups • similar controls & relationships as for volcano-bounded basins (windward side, complex chronostratigraphy, etc.) • extensional deformation common; post-rift thermal subsidence will affect stratigraphic development • internal growth strata possible if platforms on actively deforming fault-bounded highs substrates: • volcanics, volcaniclastics, faulted arc basement or volcanics/volcaniclastics • platforms become larger & thicker as depocenters fill or as deformation wanes
Fault-bounded intra-arc basins
Figure 15. Types of carbonate platforms and buildups within intra-arc basins. Platform thicknesses, widths, and relief are not to scale. Only generalized platform types and buildups are shown. Isolated platforms and fringing platforms-reefs are most common. Stratigraphic relationships and depositional contacts with structural features and various geomorphologic surfaces across intra-arc basins are also highly generalized. Modified from Smith and Landis (1995) to emphasize carbonate-platform facies.
Extensional backarc basins are commonly associated with subduction zones where trench rollback occurs. Many extensional backarc basins actually begin as fault-bounded intra-arc basins, when a volcanic arc begins to rift apart. An extensional backarc basin might come into existence when there is the first evidence for a topographically expressed, but volcanically extinct, remnant arc that is separated from an active volcanic arc by a rift zone. There likely will be a distinct bathymetric or topographic axis to the newly formed extensional backarc basin at this point in its tectonic evolution. This definition for an extensional backarc basin does not depend on the amount of stretching across the rift zone or on whether rifting has ceased or seafloor spreading has begun. The remnant arc also can be both a source area for siliciclastic sediment and a substrate for shallow-marine carbonate sedimentation, although carbonate facies may be deposited only during early stages of the basin’s history, before the remnant arc has undergone significant thermal subsidence. Backarc basins that are underlain by trapped fragments of older oceanic crust have different tectonic and subsidence histories than those that form by backarc extension. Remnant-ocean backarc basins are associated with intraoceanic subduction zones, and their dimensions depend on where subduction is initiated within the overriding plate and the nearest bathymetric high
(commonly a continental margin) that defines the distal side of the basin (see below). Backarc regions typically are areas of overall crustal extension or strike-slip deformation. Crustal shortening is rare in most backarc regions except for some continental-margin arc systems where the subducting and overriding plates are strongly coupled (e.g., segments of the Andean and western Indonesian backarc regions and the Sea of Japan backarc region). In cases where significant crustal shortening occurs in the backarc region of continental-margin arc systems, the flexural basins that develop are more appropriately described as retro-arc foreland basins, which are not considered further in this paper. The styles and intensity of deformation in the backarc region depend largely on (1) the tectono-evolutionary stage (or maturity) of the arc system; (2) the relative motions (e.g., angle of convergence) between the subducting and overriding plates; (3) the age, thickness, dip, and crustal type of the subducting slab; (4) possible changes in subduction polarity during convergence; (5) whether trench rollback is possible; (6) whether a spreading ridge, seamount, or other large-scale seafloor features have been subducted; or (7) whether arc-continent or continent-continent collision has begun. Extensional backarc regions are dominated by normal-fault networks that are subparallel or slightly oblique to the trend of
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Dorobek old oceanic crust trapped by intraplate subduction zone
A 0
100 km
new oceanic crust created by back-arc spreading
B 0
100 km
continental crust beneath retro-arc foreland basin (compressional arc system)
C 0
D
200 km
rifted continental crust beneath extensional back-arc basin
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Figure 16. Cross sections through different types of convergent margins, showing crustal types of backarc regions. (A) Intraoceanic volcanic arc, with old oceanic crust trapped in backarc region. (B) Backarc region underlain by newly created oceanic crust created by spreading axis. (C) Compressional backarc region (retro-arc foreland basin system) underlain by continental crust. (D) Extensional backarc region underlain by rifted continental crust.
the volcanic arc. As in other extensional settings, oblique-slip transfer faults trend at high angles to basin-bounding normal faults and accommodate displacement transfer between offset or nonparallel normal faults. This style of extensional deformation is characteristic of the 700-km-long Izu-Bonin intraoceanic arc system, which is considered to be an example of an extensional arc system where an incipient backarc basin(s) is forming (Taylor et al., 1990, 1991; Klaus et al., 1992). The extensional sub-basins along the Isu-Bonin arc are typically asymmetric, 25–110-kmlong and 25–40-km-wide features. Most basin-bounding normal faults trend subparallel to the axis of the arc system. Similar to continental rift systems, oblique transfer faults and accommodation zones link oppositely dipping normal faults that bound the rift depocenters. Submarine volcanic centers appear to be
concentrated along bends in the rift system that are interpreted as accommodation zones. Modern examples of backarc extension behind continental-margin volcanic arcs include the Andaman Sea, the Okinawa Trough, and the Japan Sea. Highly oblique subduction from westernmost Indonesia northward to the continental margin of Myanmar, strong coupling between the subducting Indian-Australian plate and overriding Eurasian plate, and subduction of an aseismic ridge may explain the backarc rifting and seafloor spreading in the Andaman Sea (Eguchi et al., 1980). The Okinawa Trough is an incipient continental backarc basin and displays similar structural features as the Sumisu rift zone (Letouzey and Kimura, 1986). The Japan Sea is a fully developed continental-margin backarc basin, where contraction or transpression has occurred along its eastern side, and subduction may be in its beginning stages (Kikuchi et al., 1991; Tamaki, 1995; Okamura et al., 1995; Takano, 2002). Retro-arc foreland basins are flexural depocenters that result from crustal shortening on the continental side of the volcanic arc; examples include the Andean and western Indonesian “backarc” settings. In comparison, crustal shortening within backarc basins (sensu stricto) along intraoceanic convergent margins is apparently rare. The limited amount of shortening in the backarc region of most intraoceanic volcanic arcs probably reflects weak coupling between the subducting and overriding plates along these types of convergent margins. Platform Types The lateral extent and cumulative thickness of carbonate facies within backarc basins will vary, depending on the tectonic origin(s) and evolutionary stage of the basin (Fig. 17). For backarc basins that form by entrapment of old oceanic crust behind the arc, carbonate facies may be entirely limited to shallow submerged areas on the backarc side of the arc massif, and there may be no other shallow-water substrates for carbonate platforms across the backarc region. Most emergent or shallow-water substrates in these backarc basins will be constructed by volcanic activity. Continued growth of carbonate platforms and reefs in these backarc settings may be frequently interrupted by volcanic eruptions or mass-wasting events so that reefs and other carbonate facies must be reestablished after a catastrophic event destroys or buries them (Eldredge and Kropp, 1985). In contrast, extensional backarc basins will have bathymetric profiles that reflect stretching of the arc massif or other types of backarc lithosphere. During initial rifting of the volcanic arc, carbonate sequences may be deposited in intra-arc basins (see above). Where stretching continues to the point of whole lithosphere failure and a seafloor-spreading ridge develops, a backarc basin is clearly identifiable, and trailing rifted margins develop on either side of the basin. There is no consensus, however, as to when an intra-arc basin structurally, bathymetrically, or sedimentologically transitions into a backarc basin. Apparently there are no preserved examples of thick, longlived carbonate platforms within backarc basins. Backarc basins
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INTRA-OCEANIC BACKARC SETTINGS
A
shelf/ramp sequence constructed on older volcaniclastic strata
B
COMPRESSIONAL BACKARC SETTINGS
smaller isolated ramp sequence ramp sequence platforms, shoals, and reefs constructed on older constructed on foreland on antiformal highs above foreland basin strata side of basin back-arc thrust belt (commonly post-orogenic)
pinnacle reef or outer reef tract may be fault-controlled
TRAPPED OCEANIC CRUST 0
0
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C
EXTENSIONAL BACKARC SETTINGS: C. Backarc intracontinental rifts D. Backarc intracontinental rifts leading to spreading and conjugate passive margins E. Intra-arc rifting leading to spreading
200 km
RETRO-ARC FORELAND BASIN
more regional ramp/shelf sequences during post-rift stages; flexural onlap and post-rift thermal subsidence; substrates include post-rift fill and continental basement
may contain syn-rift platforms on fault-bounded highs if depocenters are marine
EXTENSIONAL BACK-ARC BASINS
D
may contain syn-rift platforms on fault-bounded highs if depocenters are marine, but more extensive ramps/shelves form during post-rift phase
thermally subsiding conjugate passive margin; similar controls on regional platform development as passive margins
may contain syn-rift platforms on fault-bounded highs if depocenters are marine, but more extensive ramps/shelves during post-rift phase
200 km
thermally subsiding remnant volcanic arc provides substrate for short-lived isolated platforms and reefs
BACK-ARC SPREADING RIDGE
BACK-ARC SPREADING RIDGE
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E
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Figure 17. Different types of backarc settings and regional carbonate platforms that may form in them. In all examples shown, the subduction zone and forearc region (not shown) are left of the volcanic arc (indicated by triangle). Horizontal scales are generalized; no vertical scale implied. Only general patterns of platform development are shown; internal stratigraphy and details about platform morphology will vary significantly in each backarc setting. (A) Intraoceanic volcanic arc, with old oceanic crust trapped in backarc region. Carbonate platforms and reefs can form on the backarc side of intraoceanic arcs. In general, substrates for these backarc carbonate systems will be constructional volcanic edifices, eroded arc-massif basement, or volcanic-volcaniclastic deposits. Local fault-controlled bathymetry may control platform dimensions and facies patterns. (B) Retro-arc foreland-basin system. Extensive carbonate sedimentation can occur in retro-arc foreland basins if they are underfilled and the continental foreland is below sea level. Thus, carbonate sedimentation is probably more common during underfilled stages of retro-arc foreland-basin development. Carbonate ramps are common, especially along the distal (foreland) side of the basin. Various styles and intensities of deformation across the foreland can significantly affect carbonate sedimentation. Foreland-basin carbonate systems are discussed in more detail in later sections. (C) Extensional backarc region underlain by rifted continental crust. If rifted continental basement is flooded early, synrift carbonate platforms may form on fault-bounded basement highs. More extensive platforms may develop during late post-rift stages, when the extensional backarc basin is nearly filled and thermal subsidence dominates. (D) Backarc region underlain by newly created oceanic crust created by spreading axis. Where stretching in the backarc region progresses to the point of whole lithosphere failure, a seafloor spreading axis will develop. If stretching began within an intra-arc setting, the trailing remnant arc becomes extinct and undergoes relatively rapid thermal subsidence. The remnant arc can serve as a substrate for carbonate-platform development, although these are usually short-lived platforms. (E) If stretching began within an intra-arc setting, conjugate passive margins will develop on both sides of the newly created backarc ocean basin. Both conjugate margins should behave like other passive margins. Thus, carbonate platforms and reefs should become more areally extensive over time as the margins undergo thermal subsidence and progressively wider potential substrates become available for carbonate sedimentation.
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Dorobek
that are underlain by trapped oceanic crust are deep-water settings almost from their inception, so shallow-marine carbonate facies will not be deposited until the volcanic arc shoals to shallow-water conditions. Even after the volcanic arc has built itself up to shallow-water depths, backarc carbonate facies generally will consist only of fringing reefs or narrow rimmed-shelf platforms that are constructed on the steep-sided volcanic edifice. More areally extensive carbonate platforms can be constructed only after appropriate substrates for shallow-marine carbonate sedimentation have been constructed, which requires some combination of erosion of the volcanic edifice (by both subaerial and submarine processes) and deposition of volcaniclastic sediment and lava flows. Similarly, most carbonate sequences within extensional backarc basins consist of fringing reefs or relatively narrow ramps and rimmed shelves that are constructed (1) on the backarc side of the active volcanic arc and arc massif, (2) for a short time on the remnant arc as it drifts away from the active arc, or (3) on fault-bounded basement highs that partition extensional backarc regions into sub-basins. Postrift thermal subsidence ultimately submerges the drifting remnant arc to subphotic water depths so that any productive carbonate factories on the remnant-arc crust are soon terminated. Scattered pinnacle reefs and mounded buildups that are typically less than a few hundred meters thick may be constructed on synrift topography in backarc regions. These buildups apparently are drowned soon after they form because of the combined effects of rapid subsidence and short-term eustatic sea-level rise. Compressional backarc settings are more like retro-arc foreland basins, so they are not discussed here. Regardless, carbonate platforms and reefs are common in these settings, forming on fault-bounded highs that result from inversion tectonics. Isolated platforms and buildups are common morphologies. Quaternary examples. Quaternary carbonate platforms and reefs are found in the backarc basins of the eastern Indonesian archipelago, the Marianas, and Fiji. Quaternary carbonate reefs and platforms from modern backarc settings may overlap, either geographically or conceptually, with intra-arc carbonate systems, especially where fault-bounded intra-arc basins evolve into backarc basins during progressive stretching of a volcanic arc. Most Quaternary carbonate systems in backarc basins consist of (1) predominantly fringing reefs or barrier reef tracts with narrow back-reef “lagoons” that are constructed on the backarc side of the volcanic arc, (2) drowned pinnacle reefs or mounded carbonate buildups that also typically form on fault-bounded highs in the backarc region, or (3) relatively narrow platforms that are constructed on faulted backarc basement highs that either extend from the volcanic arc or serve as accommodation zones that partition extensional backarc regions into separate sub-basins. The present-day abundance of fringing reef tracts and drowned buildups reflects both the high-amplitude (i.e., >100 m) character of the post-Pleistocene sea-level rise and the typically steep structural or volcanic topography that was flooded by this rise in sea level. The obvious structural controls on platform dimensions and
locations reflect the ongoing faulting and underfilled character of many Quaternary backarc basins, which typically are in early stages of development (i.e.,
