IGC FIELD TRIP T311ORIGIN OF THE MIOCENE MONTEREY FORMATION IN CALIFORNIA

by Thomas C. MacKinnonChevron U.S.A., San Ramon, California

INTRODUCTION

Diatomaceous rocks and their diageneticequivalents, chert, porcelanite andsiliceous mudstone are abundant in Miocenedeposits of the Pacific region. Of these,the Monterey Formation is the best known andmost extensive. It is present throughoutlarge parts of the state of California(Figure 1) and is in places over threekilometers thick. In addition, it includesthe largest diatomite quarry in the worldand more importantly, is the source andreservoir for much of the oil produced inCalifornia (Figure 2).

The origin of Miocene siliceous depositsin the Pacific region is tied to afortuitous combination of tectonic, climaticand oceanographic conditions that promotedbiogenic productivity while reducingterrigenous input to coastal basins. Thisresulted in rapid accumulation of relativelyundiluted biogenic sediment, not onlysiliceous but calcareous as well, during theMiocene times.

Early work on the Monterey was dominatedby Bramlette (1946) who documented theextent of the formation and recognized itsbasic origin. This work still remains asthe most detailed statewide stratigraphicsummary.

Following Bramlette's work, surprisinglylittle work was done on the MontereyFormation until the early 1980's when aflood of publications began to appear. Theincreased interest was due in part to newanalytical methods that were capable ofextracting some of the detailedpaleoclimatic and oceanographic informationthat some Monterey Formation rocks contain,and in part due to a strong surge inpetroleum exploration for Monterey Formationfractured reservoirs. Much of the new workis included in symposium or guidebookvolumes edited by Garrison and Douglas(1981), Issacs (1981a), Williams and Graham(1982), Isaacs and Garrison (1983), Garrisonand others (1984), and Casey and Barron(1986). More recent work with up to datereference lists include: Graham and Williams(1985) on the San Joaquin Valley; Barron(1986) on dating and paleoceanography;

T311 :

Garrison and others (1987 and in press) onphosphatic rocks and paleoceanography;Isaacs and Petersen (1987) on petroleumgeology; and Snyder (1987) on structure.

GEOLOGIC SETTING

California has been part of a subductionor transform plate margin since the LatePaleozoic. Subduction appears to have beenthe dominant mode at least from the LateJurassic to Mid Tertiary times. In LateOligocene times (ca 29 Ma) subduction beganto cease due to the collision of aspreading ridge with the California margintrench (Atwater, 1970).

The change from a subduction to atransform margin resulted in a dramaticchange of depositional patterns inCalifornia that helped set the stage forMonterey Formation deposition. In the LateOligocene-Early Miocene, the Californiamargin experienced rapid downdropping andthe development of deep marine coastalbasins (Tennyson, 1986; Bachman, 1988).This event is documented by a transgressivesequence (non-marine to shallow~marine todeep-marine) that is present throughout theCoast and Transverse Ranges of California(Pisciotto and Garrison, 1981; Ingle, 1981).It is not clear whether the initialdowndropping was related wholly to transformtectonics or to some aspect of the trench­transform transftion. Regardless, extensionwas the dominant mechanism as evidenced bynormal faulting and vulcanism which occurredthroughout much of California in Early

'Miocene times (Yeats, 1987).Clastic sedimentation dominated the early

phases of fill in the newly formed basins.Then approximately 17.5 Ma, Montereydeposition began abruptly throughout much ofCalifornia. The major controlling factorwas an abrupt and major reduction in thesupply of terrigenous material to reach theoffshore areas and a high biogenicsedimentation rate (Isaacs, 1985; Graham andWilliams, 1985).

The sudden reduction in sediment supplywas apparently caused by tectonic" andpossibly eustatic events (Figure 3). A

major tectonic reorganization of theborderland, at least in Southern California,apparently occurred around the onset ofMonterey deposition. Evidence for thisincludes the beginning of rotation of partsof the Transverse Ranges and Mojave area,clear evidence of transform faulting, rapiddowndropping and uplift in coastal areas,and local volcanism, all occurringapproximately 16 to 18 Ma (Tennyson, 1986,Crowell 1987; Luyendyk and Hornafius, 1987).By Middle Miocene times, the offshore areawas probably composed of a number of small,deep (1000-2000 m; 3300-6600 ft) basinssimilar to the Southern Californiaborderland today. Outer basins wereprotected from sediment influx by trappingof coarse clastics in inner basins.Onshore, considerable sediment was trappedin newly formed interior basins (Figure 1).In addition, a high eustatic sea levelduring the Early and Middle Miocene may havecaused flooding of coastal areas, therebyexpanding shelf areas where additionalsediment could be trapped.

In addition to reduced sediment supply,high productivity of siliceous organisms inthe Pacific region was an important factorpromoting Monterey deposition. Worldwidesiliceous sedimentation apparently switchedfrom low latitude North Atlantic areas tothe North Pacific approximately 17-18 Ma dueto a change in ocean circulation (Keller andBarron, 1983). Furthermore, major polarcooling began about 15 Ma (Woodruff andothers, 1981); this probably resulted inincreased ocean circulation and upwellingwhich enhanced productivity, particularly ofdiatoms. Average silica accumulation ratesduring the Miocene are comparable withpresent rates in highly productive oceanicareas (Isaacs, 1985). High oceanproductivity and low terrigenous inputcontinued into the Late Miocene resulting inthick accumulation of hemipelagic calcareous

Figure 1. Present generalized location ofLate Miocene marine and nonmarine depositsin California. Modified from Reed (1933)and Addicott (1968).C = Coalinga; B = Bakersfield

100 Miles

tN

o 100 200 Km, i !

o 100 Miles

SALINAS ANDCENTRAL COAST

(0.3)

Figure 2. Location of major onshore Neogenedepocenters (basins) in California. TheMonterey Formation is present in parts orall of each basin (compare Figure 1).Numbers indicate cumulative oil productionin billions of barrels; approximately 90%was produced from either the MontereyFormation or clastic rocks of Middle Miocenethrough Pliocene age (Taylor, 1976). TheMonterey Formation is the source of most ofthe oil.

MID TO LATEMIOCENE DEPOSITS

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200Km!

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and siliceous deposits throughout much ofCalifornia.

Approximately 6 million years ago, asudden influx of terrigenous materialsignaled the end of Monterey deposition inmost areas of California. This influx wasmainly due to the onset of compression and

mountain building throughout most of coastalCalifornia and may have also been influencedby falling sea level. The Transverse Rangesand part of the Coast Ranges were initiallyuplifted at this time, while coastal basinscontinued to subside and fill, setting thescene for Monterey hydrocarbon generationand entrapment.

Figure 5. Thin-bedded sequence of chert,porcelanite, mudstone, and dolostone;siliceous facies, coastal California.

STRATIGRAPHY

The Monterey Formation was originallylaid down as diatomaceous and calcareousoozes with variable detrital content.Subsequently, most of these sediments wereburied deep enough so that the diatomaceousoozes were converted, through diagenesis,into the familiar siliceous rock types -

Figure 4. Diatomaceous sequence, coastalCalifornia.

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Figure 3. Generalized stratigraphic columnof the Monterey Formation whereuninterrupted by coarse clastics. Climaticinterpretation from Woodruff and others(1981); sea level curve from Haq and others(1988); tectonic events referenced in text;ages from Isaacs (1983), Barron (1986), andDumont (pers. comrn.).

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TABLE 1 Common Monterey Formation Rock Types

CHERT: An aphanitic siliceous rock (either opal CT or quartz) with a glassy, vitreous,or waxy luster; hardness >5; a knife will not scratch it; fractures concoidally orinto angular fragments; color varies in shades of white, black, brown, to nearlycolorless.

PORCELANITE: An aphanitic siliceous rock, (either quartz or opal CT) with a matte lusterresembling unglazed porcelain; commonly laminated; hardness <5; a knife will scratchit; fractures concoidally or into splintery fragments; light colored when weathered­brown to gray when fresh.

MUDSTONE OR SHALE (SILICEOUS): A fine-grained, massive to crudely laminated rock whichdents on impact, fractures irregularly or along bedding, and shows a waxy impressionwhen scratched.

DOLOSTONE OR MARL: Dolostone: an aphanitic to sucrosic rock which, whenpowdered, fizzes vigorously in dilute HC1; usually gray to brown when fresh-weathersto a yellow color; scratches easily but very tough, rings when hit with a hammer.Marl: A fine-grained, light to dark-colored rock that fizzes vigorously in HCl and issimilar in appearance to mudstone or shale as described above.

DIATOMACEOUS ROCK: A massive to laminated soft rock that is noticeably lightweight andvery porous; fresh samples are brownish-weathered samples are chalk white; detritus­rich varieties show a waxy impression when scratched; diatom-rich varieties powderwhen scratched.

chert, porcelanite, and siliceous mudstone(Figures 4 and 5), and calcareous-oozesbecame calcareous or dolomitic mudstone ordolostone (Table 1). At present,diatomaceous rocks remain only in area~

where burial was minimal, such as alongbasin margins or at the stratigraphic top ofthe Monterey Formation.

The thickness of the Monterey Formationis highly variable. The thickest depositsare found in basin centers, such as in thesouthwest part of the San Joaquin basinwhere the Monterey is over 3000 m (9800 ft)thick. More typically the formation rangesfrom 300 to 1000 m (980-3300 ft) thick(Figure 6) with thinner depositspreferentially located on what werepersistent bathymetric highs and basinmargins.

The distinctive, thin-bedded biogeniccharacter of the Monterey Formation is bestdeveloped in areas that were most protectedfrom terrigenous input from land. Ingeneral, the purest biogenic sections arefound in the Coast Ranges and the westernTransverse Ranges, areas that were probablylocated furthest from land in Miocene times.Clastic-dominated sections coeval withbiogenic Monterey deposits, are typicallyfound in easternward outcrops or in areasthat had access to turbidites.

An example of distal to proximal faciesrelationships in southern California isshown in Figure 6. In this area, theMonterey changes from nearly purehemipelagic deposits near Point Conception ­to subequal clastics in the coeval ModeloFormation in the Santa Monica Mountains - to

dominantly clastics in the Puente Formationin the Puente Hills.

Examples of turbidite sandstone bodieswithin the Monterey Formation include thehighly oil productive Stevens sandstone inthe San Joaquin Valley (Graham and Williams,1985) and lower Monterey sandstone in thewestern Santa Barbara Channel offshore (Ogleand others, 1987).

The base of the Monterey, where notobscured by coarse clastics, is generallysharp and dates at approximately 17.5 Ma(Barron, 1986). The upper contact of theMonterey Formation, where not erosional, issharp in some areas and apparentlygradational in others. Where the contact isgradational, overlying strata are siliceousbut are generally more detritus-rich thanunderlying Monterey rocks. Age of theyoungest Monterey Formation rocks isapproximately 6 Ma (Dumont and others, 1986;Barron, 1986).

In many areas of California the MontereyFormation can be subdivided into threefacies: a lower calcareous facies; a middlephosphatic facies; and an upper siliceousfacies. Additional subdivisions arerecognized locally (e.g., Figure 6).

The CALCAREOUS FACIES consists mainly ofmassive to laminated dolostone, andcalcareous or dolomitic mudstone, siliceousshale, porcelanite, and rare chert.Sections are generally thin-bedded butlithologic variation is wide with so~esections composed mainly of carbonate-richshale, others dominated by dolostone, andstill others containing abundant carbonate­rich siliceous mudstone and porcelanite.

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The PHOSPHATIC FACIES is characterized bydark, laminated, organic-rich, phosphatic,calcareous or dolomitic mudstone or marl.Dolostone nodules or beds are typicallypresent. The phosphate most commonly occursin light-colored laminations or in blebs upto a few centimeters long.

The SILICEOUS FACIES is characterized bythe predominance of laminated siliceous rocktypes - mainly porcelanite, siliceousmudstone, and chert. In addition,dolostone, and carbonate-rich mudstone arecommon in many areas. Two distinct end­member outcrop types are recognizable. Onone extreme are outcrops that are distinctlyheterogeneous with widely varying rock

types, abundant chert and dolostone, andhighly variable bed thickness. On the otherextreme are outcrops composed almostentirely of uniformly thin-bedded siliceousmudstone and porcelanite.

The three facies are not presenteverywhere and are partly age controlled(Bramlette, 1946; Pisciotto and Garrison,1981). The siliceous facies is the mostdistinctive and widespread. Its age spansthe age range of the Monterey formation(17.5 to 6 Ma), but it is best developed lnmost areas starting around 12-13 Ma ago.The calcareous and phosphatic facies arebest developed in the Central Coast Rangesand Western Transverse Ranges; they are

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Figure 6. Comparison of five Miocene sections in Southern California. The MontereyFormation in the western three sections contain easily recognizable calcareous, phosphaticand siliceous facies (shaded). Local terminology shown by letters: CS = clayey-sil{ceousmember; UCS = upper calcareous-siliceous member; P = phosphatic member; Les = lowercalcareous-siliceous member. Note that the two e.astern sections (paleo-inboard) aredominantly sandy but still contain siliceous rocks, especially of Mohnian age. Age controlis from references cited in Figure and M. P. Dumont (pers. comm.).

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'ROCK CLASSIFICATION

Hand specimen characteristics ofdiagenetica11y altered Monterey rocks (Table1) are in large part controlled by therelative proportion of three main

Figure 7. Diatomaceous rocks (opal A) aretransformed to opal CT at approximately 45°to 50° C (113° to 122° F) and to quartz atapproximately 75° to 85° C (167° to 185° F)(Isaacs, 1987). Paleo-geothermal gxadientswill control the depth of thesetransformations. In this diagram, theentire section is imagined to bediatomaceous and typical present-daygeothermal gradients are used.

1km.(3281')40°C(104°F)

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SEA LEVEL, 15°c(59°F)

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2km.(6562')115°C(242°F)

the stratigraphic thickness of both the opalA to opal CT and opal CT to quartztransition intervals varies from 20 to 300 m(66 to 980 ft) (Keller and Isaacs 1985).

Because temperature is the main control,burial depth required for silicatransformations depends on local heat flow.The geothermal gradient typically variesfrom approximately 25° to 50° C/km (1.4° to2.7° F/100 ft) in Monterey Formation basinsin California. This translates into aburial depth range of 600 to 1400 m (1970 to4590 ft) for the opal A to opal CTtransition and 1200 to 2800 m (3940 to 9190ft) for the opal CT to quartz transition,for onshore sections {using surfacetemperature = 15°C, (59° F) and assuming aconstant past geothermal gradient; Figure7). These calculations are consistent withwell data.

SILICEOUS ROCK TYPES AND DIAGENESIS

poorly represented or absent in many otherareas, most notably in parts of the SanJoaquin Valley. Where present, thecalcareous facies typically extends from17.5 Ma (base of Monterey) to 13 to 14 Ma;the phosphatic member most commonly rangesfrom 12-15 Ma (Figure 3).

The shells of diatoms, radiolaria andother siliceous organisms are composed ofopal A, a hydrous silica phase (Si02.nH20)that is inherently unstable. With burial,and mainly due to increasing temperature,opal A dehydrates and converts to a silicaphase similar to crystoba1ite-tridymitecalled opal CT (Si02). With a furtherincrease in burial, opal CT converts toordinary quartz.

The transformation from opal A to opal CTto quartz results in major changes in rockproperties and appearance (Murata andLarson, 1975; Isaacs, 1981a, 1981b). Themost dramatic change occurs in thetransformation of soft, punky diatomaceousrocks (opal A) to chert, porce1anite orsiliceous mudstone (opal CT) which aretypically hard and brittle. The change IS

accompanied by a sharp drop in porosity(Isaacs, 1981b) and an increase in grain andbulk density.

A further decrease in porosity anddensity values occurs during thetransformation of opal CT to quartz.However, in this transformation, the handspecimen appearance of rocks does notusually change appreciably. For example, anopal CT chert normally cannot bedistinguished from a quartz chert in handspecimen. Generally, the only sure way todistinguish opal CT and quartz phase rocksis by x-ray diffraction, though in somecases, particularly with cherts, thinsection examination may suffice.

Temperature appears to be the majorfactor controlling silica phase transfor­mations. Various studies (summarized inIsaacs, 1987) have shown that thetransformation temperatures for opal A to CTare approximately 45° to 50° C (113° to 122°F) and for opal CT to quartz areapproximately 75° to 85° C (167° to 185° F).In addition, composition has a minor butimportant control on transformationtemperature - rocks with highsilica/detritus ratios convert from opal Ato opal CT at lower temperatures and from CTto quartz at higher temperatures than rockswith low silica/detritus ratios (Isaacs,1982). Outcrop and well studies show that

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components: (1) biogenic silica; (2)detritus; and (3) carbonate (Isaacs, 1981a).For example, geochemical analysisconsistently show that cherts containgreater than four parts biogenic silica toone part detritus. As another example,analyses of typical porcelanites givesilica-detritus ratios of from 1:1 to 1:4.Similarly, siliceous mudstones anddolostones contain high percentages ofdetritus and carbonate respectively. Suchdata provide a means of classifyingsiliceous rocks quantitatively.

The proportion of biogenic silica,detritus, and carbonate is determined usinga geochemical method such as described byCarpenter (this volume). In this procedure,the elemental composition of siliceoussamples is determined by x-ray fluorescenceand other methods. The elemental data arethen converted to mineral composition usingformulas based on detailed geochemicalanalyses of control samples from nearbywells or outcrops. Finally, the proportionsof biogenic silica, detritus and carbonateare calculated from the mineralogy.

The proportions of biogenic silica,detritus, and carbonate can conveniently beplotted on a triangular diagram to serve asa quantitative basis for classification.Our work shows that cherts, porcelanites,shale/mudstone, and marlstone/dolostone, as

identified by field characteristics,consistently fall in the broadly definedcompositional fields shown in Figure 8.

Isaacs (1982) pointed out potentialproblems with a compositional-based schemefor rock classification. She noted that thetextures of chert, porcelanite and siliceousmudstone result in part from thecrystallization character of silica. Shenoted that clay type, rather than simply theamount, might affect crystallizationcharacter. She further noted that coarsesilt and sand-sized detrital grains are lesslikely to influence silica crystallinecharacter than an equal volume of dispersedclay-sized particles. We have not noted anycorrelation of clay type to texture,however, it is obvious that the presence ofcoarser detrital material in an otherwisefine-grained rock could give erroneousresults. Therefore, bimodal rocks such assandy diatomites and their diageneticequivalents may require specialconsideration. Regardless of theseproblems, our compositional-based scheme(Figure 8) provides a useful basis forquantifying and comparing field-baseddefinitions.

OCEANOGRAPHIC, CLIMATIC, AND TECTONICCONTROLS ON SEDIMENTATION

DETRITUS The complex interbedding of differentrock types that characterizes the MontereyFormation is the result of variations in theamount of detritus, biogenic silica andbiogenic carbonate deposited at a giventime. The relative abundance of these threecomponents depends on a complex interplay ofclimatic, oceanographic and tectonicconditions (Garrison, 1981; Kennett, 1982).

Detrital input in the marine environmentis largely from fluvial sources (Gorslineand Douglas, 1987) and is stronglyinfluenced by sea level changes, tectonics,and climate. For example, sediment supplywill increase if sea level falls, uplift of'mountains is initiated or accelerated, andgenerally in periods of higher rainfall orfloods. These factors must have played animportant role in controlling sediment inputinto Monterey Formation basins.

Biogenic silica and carbonate consistmainly of microscopic,shell-forming plantsand animals. The plants live in the photiczone, or upper 50 to 100 m (160 to 330 ft)of the water column, and consist mainly ofdiatoms (siliceous) and coccolithophores(calcareous). The animals consist mainly offoraminifera (calcareous) and radiolaria(siliceous); these live anywhere in the

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Figure 8. Generalized geochemicalclassification of Monterey rock types. Rockidentification from field characteristics IS

somewhat operator dependent and thereforethe boundaries between rock types isapproximate.

BIOGENICSILICA

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water column or on the ocean floor, but manychoose to live in the photic zone because ofhigh nutrient levels there.

Total productivity of the biogenicpopulation is controlled by the supply ofnutrients available. Areas of highproductivity are generally associated withupwelling zones which form when winds orocean currents displace surface waters,allowing underlying waters to upwell (Figure9). Upwelling waters bring nutrients,mainly phosphorus and nitrogen derived fromthe decay of biogenic organisms, from depthsof a few hundred meters to the surfacewaters where they can be reutilized. Undersuch conditions, total productivity can bean order of magnitude greater than in areaswhere upwelling is not present.

An important feature of some upwellingareas is an underlying layer of oxygen­depleted water known as the oxygen minimumzone. These zones are created by oxygendepletion due to the decay and oxidation ofdead organisms as they fall through thewater column. Above this zone, surfacemixing provides oxygen; below this zone,there are fewer organisms present to use upavailiable oxygen by respiration or decay.Well developed oxygen minimum zones range indepth from roughly 150 to 1500 m (490 to4900 ft) (Figure 9). Where an oxygenminimum zone impinges on the ocean floor,little oxidation can take place and organicmatter is more likely to be preserved. Inaddition, large burrowing organisms cannotsurvive, and thin beds or laminations arepreserved rather than destroyed bybioturbation.

Of particular importance to the MontereyFormation is the observation that diatomsout compete all other shelled microorganisms

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Figure 9. Model for hemipelagicsedimentation in a nearshore upwellingregime; adapted from Rullkotter (1983).

in oceanic areas of high nutrientconcentration associated with upwellingareas (Garrison, 1981). In areas of low tomoderate productivity, calcareous organismsgenerally predominate.

Preservation of shell material is also animportant variable in determining sedimentcomposition (Garrison, 1981; Kennett, 1982).Calcareous shells are subject to increaseddissolution with depth. The depth at whichcalcareous shells cannot survive (calciumcarbonate compensation depth-CCD) averagesapproximately 4500 m (15,000 ft) in theworld's oceans. However, the CCD risestowards the continents and may possibly bevery shallow locally, particularly in areasof high productivity where decay of abundantorganic matter generates C02 which increasescarbonate dissolution. Siliceous shellscomposed of unstable opal A are also subjectto dissolution. In many areas of lowproductivity, siliceous shell material isabsent in sediments because dissolutionoccurs before burial can take place.Protection of both siliceous and calcareousmaterial from dissolution is enhanced byincorporation into fecal pellets.

Cycles

Because the input of biogenic anddetrital material is dependent on so manyvariables, frequent variations should beexpected. The thin-bedded, laminated andunlaminated rocks of the Monterey Formationattest to this and provide a detailed recordof paleoceanographic events. Some of thesechanges were cyclic in nature as evidencedby the presence of laminations and cyclicbedding in the Monterey Formation(Bramlette, 1946; Pisciotto and Garrison,1981; Isaacs, 1983).

Laminations in the Monterey Formation areexamples of clastic-biogenic cycles ofclimatic origin. Most of the laminationsare apparently varves reflecting seasonalchanges in sediment supply (Bramlette,1946). Dark layers are typically detritus­rich and represent periods of high rainfalland runoff on land (winter), whereas thelight-colored layers are concentrations ofrelatively undiluted biogenic siliceousmaterial (summer). Many modern analogueshave been recognized, such as oxygen-starvedareas of the Santa Barbara Channel and partsof the Gulf of California (Donegan andSchrader, 1981).

Somewhat longer term climaticfluctuations appear to have created much ofthe thin-bedded character of the MontereyFormation. For example, alternating

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detritus-rich versus biogenic-rich beds,typically from a few to 30 em (1 to 12 in)in thickness are thought to represent timeperiods of 10's to 100's of years, asdetermined from counting laminationsinferred to be varves (Pisciotto andGarrison 1981). In some cases, the clasticpart of these cycles appears to be ofturbidite origin. In other cases theclastic portion appears to representsustained rather than a sudden influx ofdetrital material. In both cases, climaticfluctuation, such as periodic floods or dry­wet cycles of longer duration may be themain controlling factor.

An additional explanation for the thinbedded character of the Monterey Formationincludes short term changes in oceanographicconditions. For example, silica-carbonatevariations may be due to preferentialpreservation or changing temperature­nutrient conditions that favor the growth ofcertain organisms. As an example, unusuallypure diatomaceous beds (chert precursors?)may have been deposited during short periodsof intense upwelling, which created diatomblooms while promoting dissolution ofcarbonate.

Alternating laminated and non-laminatedintervals represent still another type ofcycle probably controlled by oceanographicconditions. These intervals are generallyfrom a meter (3.3 ft) to tens of meters (33ft +) in thickness and are most clearlyvisible in diatomaceous' sequences. Most ofthe non-laminated intervals have clearlybeen bioturbated; they appear to representperiods of oxygenated conditions duringwhich time large burrowing organisms couldsurvive. Laminated intervals represent lowoxygen conditions and the absence ofburrowing organisms. Such periodicfluctuations in oxygen are probably due toocean circulation changes.

ORIGIN OF DOLOMITE

The origin of dolomite in the MontereyFormation has been the subject of extensiveresearch in recent years (e.g., Garrison andothers, 1984). Most of the dolomite occursin three forms: (1) nodules; (2) beds; and(3) as disseminated crystals.

Disseminated dolomite is the most commontype of occurrence in the Santa Barbara­Santa Maria area (Isaacs, 1984) and perhapsin the Monterey Formation in general. Inthin section, this· dolomite is typicallyevenly dispersed as silt and clay-sizedcrystals in mudstone, porcelanite and chert.

Dolostone nodules and beds make up only a

small portion of the Monterey Formation butthey are conspicuous because they are hardand resistant to erosion and often weatherto an unusual yellow-orange color.Dolostone nodules are generally less than ameter (3 ft) thick and up to several meters(10 ft +) in length. Dolostone beds aregenerally less than a meter (3 ft) thick butrange up to 3 m (10 ft) in thickness.

Most dolostone beds and nodulesapparently formed early, while the hostsediment was still soft and uncompacted.Field evidence for this include differentialcompaction features best displayed bynodules - bedding in the host rock commonlywraps around nodules (Figure 10) andfurthermore, laminations in nodules aretypically further apart than in laterallyequivalent host rock. Laboratory evidencefor early formation includes oxygen isotopestudies which show that many beds andnodules apparently formed in the 20° to 30°C (68° to 87° F) range (see papers ~n

Garrison and others, 1984). Thiscorresponds to a burial depth of no morethan a few hundred meters (several 100 ft).

Modern analogs of early formed dolomitehave been recognized by ocean drilling inareas of organic-rich sediments as found,for example, off Peru and in the Gulf ofCalifornia. In these areas, nodules andbeds are forming today at very shallowburial depths and possibly on the sea flooritself. The origin of early dolomite isapparently promoted by the oxidation of or­ganic matter which results in excess C02.

Figure 10. Small dolostone nodule inphosphatic mudstone. Note how laminationsbend around nodule indicating that thenodule formed prior to full compaction ofthe host sediment.

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Furthermore, reduction of sulfate to sulfideis considered important because sulfate is amajor inhibitor of dolomite formation.Magnesium may be derived directly from seawater (see papers in Garrison and others,1984) •

An important aspect of early formeddolomite is that it helps preserve geologicfeatures that are usually destroyed in otherrock types. Dolostone beds and nodulescommonly contain; (1) detrital remnantmagnetism useful in paleomagnetic tectonicstudies; and (2) dateable siliceousmicrofossils in otherwise barren sequences.

Though early dolomite is common, oxygenisotope data indicate a later highertemperature origin for m~ch of thedisseminated dolomite and at least the outerparts of some nodules and beds. Some ofthis dolomite apparently formed by laterreplacement of calcite rather than directprecipitation. In addition, dolomite alsooccurs locally as a late stage diageneticvein filling (Redwine, 1981; Roehl, 1981)~

ORIGIN OF PHOSPHATIC BEDS

Phosphatic beds in the Monterey Formationare of particular interest because they typ­ically have very high total organic carbon(TOC) values and are excellent hydrocarbonsource rocks. In addition, phosphaticmaterial is a key environmental indicator.

Two main phosphatic facies have beenrecognized in the Monterey Formation byGarrison and others (in press). These are:(1) a phosphatic marlstone facies; and (2) apelletal-oolitic phosphorite facies.

Phosphatic marlstone is by far the mostcommon type. It is typically laminated,commonly contains 50% or more calcite ordolomite and is thought to have originallybeen deposited as a coccolith-foraminiferal­diatom mud (Figu~e 11). Fossilizedbacterial mats are commonly present and maybe a major component of the organic matter(Williams and Reimers, 1983). The phosphateis found in blebs and pellets or asdisseminated crystals along laminations.The phosphate appears to have formed priorto significant compaction as indicated byfield evidence for soft sediment deformationof phosphate nodules and the inclusion ofnodules in early formed dolostone.

Garrison and others (1981) suggest thatphosphatic marlstone formed by slowsedimentation in a low oxygen environment(oxygen minimum zone) dominated bycalcareous sedimentation. Such conditionsmay be met in a variety of enviroments such'as outer shelf, slope, basin floor, and on

Figure 11. Laminated phosphatic mudstonecontaining a layer of phosphatic nodules.

deeply submerged banks. Modern analoguescited by Garrison and others (1987) includethe upwelling areas off Peru and on the westAfrican shelf. In these two areas,phosphorites are forming today on or withina few centimeters (inches) of the seafloorand within oxygen minimum zones associatedwith upwelling.

Pelletal-oolitic phosphorites, asdescribed by Garrison and others (in press),consist of phosphatic sandstones some ofwhich contain'glauconite, ter~igenous

material, phosphatic micronodules,intraclasts, and fish fragments. Thesephosphorites are typically interbedded withsiliceous mudrocks, and in some casesprobably formed by winnowing of sand-sizedphosphatic grains by wave or current action.In other cases the phosphate may have formedon the sea floor without significantreworking. Pe1latal-001itic phosphoritesare though to have formed on shelves andbank tops; some of these deposits wereapparently redeposited in deeper water byturbidity currents.

Acknowledgements

I thank Robert E. Garrison and Michael P.Dum~nt for reviewing preliminary versions ofthis manuscript.

REFERENCES

Combined references are at end of thefollowing paper.

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