Differential Mobility of Elements in Burial Diagenesis of Siliciclastic Rocks

DIFFERENTIAL MOBILITY OF ELEMENTS IN BURIAL DIAGENESIS OF SILICICLASTIC ROCKS

ROBERT P. WINTSCH ~D CINDY M. KVALE Department of Geological Sciences, Indiana University, Bloomington, IN 47405..5055 USA

Aesrvacr: Mass-balance calculations involving smectite-fich siliciclastic mudstunes sampled from boreboles show that significant K20 addition and CaO loss correlate strongly with the progress of illitization of smectite and with the dissolution of calcite. The calculations are based on the immobility of AI203, and require a chemically homogeneous sediment at deposition, such that the shallower samples represent protolith compositions of deeper samples. Constancy of ratios of TiO2/AlzO3 are used to (1) establish that both TiO~ and Ai203 have similarly low differential mobilities, and (2) identify strafigraphie sections available in the literature that are sufficiently homogeneous to allow the calculations.

Our calculations show that most major-oxide components are immobile at the ~ 5% limit of resolution of the calculations. Exceptions are K20, which is almost universally added to modstones, and CaO, which is commonly removed from mndstones during diagenesis. The addition of up to 50% KzO correlates with percent illite in mixed-layer illite/smecfite, in- dicating that illitization ofsmectite involves fixation of K20 from diagenetic fluids. Removal of ~ 40% CaO from some sections correlates with de- creasing calcite mode and is attributed to dissolution of calcite. Together these reactions lead to a 10-15% loss of mineral volume, a volume loss that must be added to porosity loss when modeling compaction.

Illitization and carbonate dissolution may be linked, because K-fixation during illitisafian decreases the pH of the local fluid (a "reverse weathering reaction"), which enhances carbonate dissolution. Addition of up to 20% Na20 in some rocks may be caused by albitization of plagioclase, but Na20 addition is not pervasive. The apparent lack of mass transfer associated with dissolution of kaolinite and feldspars or precipitation of chlorite implies that these reactions are balanced by other reactions such that their oxide components in mudstones are conserved. Limited data from sandstone-mndstone pairs are consistent with transfer of K~O from sandstones to mudstone via rising diagenetic fluids. Except for CaO and the alkalis, diagenesis of siliciclastie mudstones appears isochemical (± 5% relative) for most major oxide components. The much larger changes recorded in some reservoir sandstones are clearly exceptions, and may not reflect changes occurring at the scale of kilometers.

INTRODUCTION

Petrographic, chemical, X-ray, and modal data from silieiclastic sediments have documented the importance of diagenetic reactions such as dissolution of feldspars and heavy minerals, illitization of smectite, albitization of plagioclase, dissolution-precipitation of carbonate and kaolinite, and precipitation of chlorite (Blatt 1992). To account for some of these mineralogical changes and replacement reactions, significant transport of aqueous SIP,, AI *~, Ca +-', Mg +2, Fe +~, K +, and Na + during diagenesis on the scale of hand samples or greater has been proposed (see discussion). However, the transfer of chemical components into or out of a rock is limited by the reactions taking place. During the dissolution stage of a reaction, all components of the reactant minerals are borne in the aqueous fluid, at which stage mineral solubilities become the dominant factor controlling component fluxes.

The identification of a net transfer of some components during diagenesis and metamorphism would be significant, because mass transfer would reflect the operation of reactions in the rock that modify its bulk composition. In carbonate rocks dolomitization is well known to modify mineral and rock compositions greatly. In rocks dominated by silicates, how-

ever, the effect of diagenetic reactions on the bulk composition of rocks is not well known, in spite of well documented depth-related mineralogical changes. Ifdiagenesis or metamorphism does modify the composition of siliciclastic sediments progressively with temperature and time, then comparisons of the compositions of ancient sedimentary rocks with young sediments (Taylor and McLennen 1985) may not be appropriate.

The goal of this paper is to identify through mass-balance calculations to what extent reactions between minerals and fluids modify the bulk compositions of siliciclastic rocks during diagenesis. If chemical components that participate in diagenetic reactions have low differential mobility, in spite of the chemically open system (on a kilometer scale; Land et al. 1987; Wintsch and Kvale 1991), then the progress of reactions would be limited by the concentrations of reactant components in the original sediment. If, on the other hand, reactants have large differential mobility (scale > m), then the progress of diagenetic reactions would be limiled by the transport of these components to the reaction site. We have used data on rock composition from samples collected mainly from oil and gas wells (Table 1) in our calculations. The resolution of this analysis is determined by the 50--500 m sample spacing of vertical sections available in the literature, and is the scale at which chemical transport can be identified.

This study has several limitations, the most important of which is the small number of drill holes from which chemically well analyzed samples arc available. Another is the need to treat younger, shallower samples as the compositional precursors of the deeper, diagenetically modified samples. A third is the lack of data on the concentrations of minor and trace elements, which would allow better identification of components of low differential mobility, upon which the calculations are anchored. In spite of these drawbacks our analysis does show enough reproducibility from well to well that at least preliminary conclusions can be drawn.

GEOCHEMICAL MASS BALANCE

In any analysis of metasomatic mass transfer, an apparent increase in the weight percent of a component can be caused by either addition of that component or by removal of other components; the latter would concentrate the less mobile components. The system of equations devel- oped by Gresens (1967) allows calculation of the changes in mineral volume accompanying diagenesis, if the densities and weight percents (gin/ 100 gm rock) of all components in parent (P) and daughter (D) samples are known. This relationship is:

lO0[f.(gD/gP)C, ° - C, P] = X, (1)

where f. is volume factor (Vu~,e,,~,/Vp~,,,,), g is grain density (not bulk density; see Brimhall and Dietrich 1987), Cj is the concentration of Com- ponent 1, and X, is the weight of Component 1 gained or lost by the daughter sample relative to the parent sample. Even with the composition of the parent material well known, the system of equations is still under- determined by one, and either the net gain or loss of at least one component or the change in volume during metasomatism must be established by independent means. Because independent information on the net transport of a component in geochemical systems is rarely available, most studies resort to assuming immobility of one of the components. In such cases the gain/loss of Component 1 in Eq 1 is set to zero. The volume factor, measuring the dilution or concentration of the immobile component due

JOURNAL OF SEDIMentaRY RESeaRCH, VOL. A64, No. 2, APRIL, 1994, P. 349-361 Copyright © 1994, SEPM (Society for Sedimentary Geology) 1073-130X/94/0A64-349/$03.00

350 ROBERT P. It 7,'\7SCH AND CINDY M. KVALE

TaBtE I.-- Summary of sources of rock chemical compositional data

Geo- thermal

Formalion Depth Range NO. of Elements nol Ca'adient | tJ3calily Reference (Basin) Age of Wells Ira) Anal.* Determinedt (Kay"C)

Alexandria Area, a van Moon 1971 Wilcox Eocene 1536--1623 2 ? Louisiana b van Moon 197 l 3062-3203 7

c van Moon 1971 3692-4837 10

South Pass, La. d" Weaver and Beck 1971 Miocene-Pliocene 1290-5014 21 Si, Ti, Fe, Ca, Na 24 e Chaudhuri and Cullers 1979 Miocene-Pliocene 1810-..4770 4 [REE only]

Harris County, Texas f" Hower eta[. 1976 Anahuac L. Miocene 1850-5500 13 30 Frio U. Oligocone

Galveston County, g' Perry and Hower 1970 Anahuac Oligocene 2141-3648 19 33 Texas Frio

Oklahoma h Weaver and Beck, 1971 Springer Mississippian 6322-7196 8 Si, Ti 23 Shale

Cameroun i' Dunoyer de Sagonzac 1964 (Douala) Cret-Paleocene 770-3900 24 K, Na, Ca 39

Kenedy County, j ' Calvert and Klimentidis, 1986 Frio Oligocene 1407-5846 66" Ti 32 Texas k' 2161-5845 22 Ti 32

SW Texas I Boles t978 Wilcox Eocene 2592-4641 5, 5 ?

Montana rn Eslinger and Sellars 1981 Belt Siers Late Proter. (8.5 km section) 26 ?

Papau Barikewa n van Moon 1971 (Kutubu Trough) Albian-L. Jurassic 1067-4232 13 ? Omati o van Moon 1971 3214-4371 7 ?

Brazil p' Chang et al. 1986 (Potiguar) Cretaceous 1212-4417 5, 5 33 q' Chang et al. 1986 (I. Santana) Cretaceous 1632-3830 5, 5 28 r" Chang et al. 1986 (Cassipora) Cretaceous 1511-3437 6, 6 22 s (Ceara) Cretaceous 1475-2095 2 32

r Muffler and White 1969 (Salton Trough) Pliocene 140-1590 10 68 u' Muter and White 1969 (Salton Trough) Pliocene 622-2871 6 ~ 300 v' Jennings and Thompson 1986 (Salton Trough) Pliocene ? 18 ? w" Yau et al. 1988 (Salton Trough) Pliocene 384-1534 7 ~250

Imperial Valley, Calif.

* Number of samples analyzed; two numbers indicate sandstone/shale pairs. t Includes H20, CO2 and ferrous iron in all cases. § Average calculated from data in reference. • Sections discussed in text. • Composite profile of six closely spaced wells.

to mobility of the other components, is then calculated and used in cal- culating the gains or losses of the other components in the daughter sample.

VOLUME CHANGES

It is possible that volume changes in a rock attdbutahle to minerals (not porosity) could lead to changes in the concentration of some elements. This might occur through dissolution or precipitation of minerals (cements) or from changes in crystal structure (e.g., aragonite ~ calcite; feldspars ~ illite). Such reactions are significant because they affect the porosity and permeability of the rock, and because changes in bulk mineral density could affect the way compaction in sedimentary sequences is mod- eled. Eq 1 is solved accurately for volume changes only when specific- gravity data for each sample are available.

SELECI'ION OF DATA

In order to assess changes in the composition of siliciclastic rocks due to progressive diagenesis, the composition of the unmodified sediment must be known. However, differences in original sediment composition caused by facies changes could be large, and the compositions of younger, shallow samples may not be good estimates of the starting composition for the deeper samples. For the present purposes only those studies that sampled a single rock type with a composition at deposition that is relatively constant with time (depth) can be considered. Because of the differences that may exist within and among different sedimentary sections, only those changes that are reprodudble in a variety of environments, rock types, and wells can be taken as representative of general diagenetic reactions.

As a test for original sediment homogeneity, we follow Gresens' (1967) suggestion of identi~ing a constant ratio of components in a related suite of samples. This is much superior to any assumptions about the low differential mobility of any single component. The rationale is that a constant ratio of two or more components in both parent and daughter samples would require that each be added or removed in proportion to the concentrations present in the parent. This is very unlikely because most components are present in more than one mineral, and proportional gains and losses would require proportional reactions of more than one mineral over a wide range of temperature, time, and depth. Thus, the alternative interpretation, that both components have low differential mobility, is much more likely. A constant ratio would further imply that the composition of the sediments at deposition was uniform. The greater the number of related samples with a constant ratio of components, the stronger the argument of immobility becomes. Similarly, the greater the number of components with a constant ratio, the stronger the argument that they all have low differential mobility. We (Wintsch et al. 1991) have success- fully applied this procedure to the problem of differential mobility during pressure solution ofmudstones of the Ordovician Maainsburg Formation on the scale of tens to hundreds of meters. Here we explore the application of mass-balance calculations to younger mudstones at a scale ten times greater.

Experimental work (e.g., Thornton and Seyfried 1985) shows that the least soluble major components in sedimentary rocks are (in increasing order ofsolubility) TIP2, A1203, Fe203, and SIP:, and thus are the most likely to have low differential mobility. The low differential mobility of these components is supported by our analysis of ~ 100 samples of Mar- tinsburg Formation mudstones in which these components were shown

DIAGE:~7~SIS OF SILICICLASTIC ROCKS

TABLr 2.--Summary of changes in component concentration with depth*

351

Dunoyer Calvert and Calvert and Weaver and Hower et al. Perry and de Segonzac Klimentidis 1986 Kl.im~tidis 1986

Camp Beck 1971 1976 Hower 1970 1 9 6 4 Well 8 Wells 1-,-.8

SiO, 0.17 (1.13) -3.95 (2.62) 10.79 (2.12) -0.65 (1.30) -2.45 (I.1 I) H 0.0021 0.1184 0.5049 0.0125 0.0699

TiO~ 0.82 (I.79) -8,52 (7.29) -9.76 (4.74) r 2 0.0185 0,0745 0.1616

FeO -5.68 (1.60) - 19,59 (3.31) 15.15 (7.00) 1.70 (5.18) -4.90(2.23) r 2 0.5334 0.6736 0.1756 0.0053 0.0695

MgO -4.38 (2.56) - 10.50 (4.82) -65.03 (5.01) 26.08 (5.05) -3.69 (I,61) - 1.95 (I.84) r: 0,1395 0.3011 0,9083 0.5479 0.2074 0.0170

CaO -204.86 (27,32) -294,15 (55.52) - 16.10 (6,91) -22.40(4.71) r-" 0.8364 0,6228 0.2133 0.2579

K~O -3.25 (1.02) 6.29 (I.18) 7.43 (4.58) 1.75 (1,49) 5.53 (I.20) r ~ 0,3589 0.7199 0,1343 0.0639 0.2466

Na~O -24.04 (5.04) -38,37 (17.14) 2.31 (6.22) 3.93 (3.08) r z 0.6739 0,2276 0.0068 0.0244

Fv -0,03 (0.02) -0.04 (0.01) -0.08 (0.02) 0.07 (0.02) -0.03 (0.01) -0.06 (O.Ol) r ~ 0.0860 0.6340 0,4725 0.3368 0.3178 0.2906

n 20 13 19 24 22 67

r, 2 (95%)i 0,197 0.306 0.208 0.163 0.179 0.06

* Values in %/kin, with standard error ( ) and calculated correlation coefficient (r2). 1- Critical values oft, z at 95% confidence, from Arkin and Colton (1963).

to be relatively immobile, even under the much more extreme metamor- phic conditions of pressure solution during development of slaty cleavage on the scale of < 1 cm (Wintsch et al. 1991). We used this "ratio test" using AI203/TiO2 to identify sections that had uniform bulk compositions at deposition because we found this method successful in our analysis of the Martinsburg Formation. Use of SiO2 in these ratios is not appropriate because important SiO2 mobility during diagenesis has been proposed (Boles and Franks 1979).

The ideal data base for this study should include a large number of samples (> 50) collected from a depth range of several kilometers, and analyzed for all major elements and many trace elements. From our search of the literature, sections a, b, c, e, h, 1, o, and s (Table 1) are of limited value to our study because of a narrow range of depths (and temperatures) sampled or because of the few elements analyzed. All rock types present were sampled in the sections presented by van Moort (1971), Boles (1978), Eslinger and SeUars (1981), and Chang et al. (1986), and so the ratio of TiO,/AI203 varies erratically. Variation in the ratio of TiO2/Al203 is small in other Gulf Coast and in all Salton Trough rocks, suggesting that not only was the composition of the original sediment relatively uniform throughout the sequence, but also the concentrations of these elements have not been modified by diagenesis.

Northern Gulf Coast.- The section with the most uniform composition is from the Oligocene and Miocene Anahuac and Frio formations from Hams County, Texas (Hower et al. 1976). These sediments were deposited during several minor, laterally coalescent, vertically repetitive deltaic cy- cles (Galloway el ah 1982). Marine processes redistributed and mixed these sediments, so the compositions of the sediments are relatively homogeneous. Hower et al. (1976) provide bulk chemical analyses of 13 samples and quantitative X-ray modal mineralogy for 14 samples to a depth of 5.5 kin. The clay minerals have also been characterized in detail by Ahn and Peacor (1985) and Lee et al. (1985). Because of its complete- ness, this data set has been reinterpreted by other workers (Boles and Franks 1979; Land, 1984). Several other studies report the bulk compositions of rocks at depth from the northern Gulf Coast region. Perry and Hower (1970) give the chemical analyses of 19 samples from Galveston County, Texas, and the mineralogy of several silt and clay size fractions of other samples; only selected samples were analyzed for both mineralogy and composition.

South Texas.-A suite of rocks with over 60 quantitative X-ray modal and partial chemical analyses deposited in a passive-margin setting in

south Texas was studied by Calvert and Klimentidis (1986). These samples are part of the Tertiary Frio Formation of the Norias delta system of south Texas. Drainage was in a semiarid, volcanic-rich region, in which sediment mixing was minimal (Galloway et al. 1982), and sediments contain variable quantities of volcanogenic and carbonate detritus. The samples come from adjacent wells in a small field in Kenedy County, and all 67 samples can be treated collectively in a composite profile. However, the 22 samples are from a single well (Table 4) that spans 1932-5227 m, and are also treated alone.

Salton Trough.- Several studies describe the compositions of rocks from the Salton Trough (Muffler and White 1969; Jennings and Thompson 1986; Yau et al. 1988). The compositions of these rocks are variable, but the parallel trends in TiO2/AI203 and other ratios of two adjacent wells (Muffler and White 1969) suggest that the compositional differences are inherited from stratigraphic differences. Because these sections have a high and variable geothermal gradient, they are discussed separately below, and related through temperature rather than depth.

In summary, almost half of mudstone sections that include bulk chem- iccal analyses (Table 1) are appropriate for the calculations in this study. These include six sections of low geothermal gradient, and four of higher geothermal gradient in the Salton Trough. Although this is a small number, the sections are sufficiently diverse that any consistent geochemical behavior among them should reveal general chemical processes during the diagenesis of smectite-rich rocks.

RESULTS OF CALCULATIONS

The above review identified both TiO2 and AI203 as immobile components during mudstone diagenesis on the scale of 0.5 km or less. This is confirmed by the small range of calculated gains and losses of TiO2 holding A1203 constant, and resultant small r 2 values (Tables 2, 3). Thus both A1203 and TiO2 are immobile during diagenesis to a - 5 % relative resolution, and either of these components could be used as an anchor for the mass-balance calculations. Unfortunately, TiOz cannot be used to anchor all these calculations because several studies did not include TiO2 in the analysis. We thus adopt A1203 to anchor the calculations below, while recognizing that further testing of this procedure with ratios of other elements, especially Zr, Y, Sc, V, Cr, and the rare-earth elements, is de- sirable.

The gains and losses of components in the I 1 suites discussed above

352 ROBERT P. WINTSCH AND CINDY M. KVALE

Ta~L~ 3.-Summary of changes in component concentration with temperature*

Mul~ler and White Mufner and White 1969 1969 Jennings and Yan et al. Salt~n Sea

Comp. Well I Well 2 Thompson 1986 1988 Composite

SiO, -0.0585 (0.041 I) -0.0630 (0.0351) 0.0060 (0.0450) 0.0952 (0.0362) -0.0257 (0.0311) r ~ 0.2017 0.4459 0.0016 0.5797 0.0198 TiO_, 0.0377 (0.0114) -0.0076 (0.0143) 0.0347 (0.0341) 0.0527 (0.0552) 0.0485 (0.0139) r" 0.5781 0.0062 0.0861 0.1541 0.2645 FeO 0.1037 (0.0435) 0.0278 (0.0246) 0.0071 (0.0315) 0.0443 (0.1276) 0.0763 (0.0241} r ~ 0.4 154 02425 0,0045 0.0235 0.2275 MgO - 0.0970 (0.0428) 0.0873 (0.0419) 0.1390 (0.0485) 0.1326 (0.1239) 0.0220 (0.0464) r 2 0.3912 0.5208 0.4276 O. 1865 0.0066 CaO -0.4670 (0.2411) 0.0665 (0.0417) -0.0414 (0.0678) -0.3141 (0.3465) -0.2302 (0.0897) r ~ 0,3192 0.3888 0.0328 0.1411 0.1624 K~O 0.1544 (0.0421) 0.1061 (0,0411) 0.1158 (0.0671) 0.2998 (0.1083) 0.1349 (0.03011) r: 0.6274 0.6248 0.2131 0.6051 0.3608 Na,O 0.1331 (0.1026} - 0.3350 (0.1527) 0,0576 (0.1033) 0.6160 (0.1129) 0.1744 (0.0755) r ' 0.1738 0.5462 0,0274 0.856 [ 0.1356 Fv -0.0008 (0.0002) -0.0003 (0.0003) 0,0002 (0.0003) 0.0008 (0.0004) -0.0002 (0.0003) r: 0.5576 0,1802 0,0500 0.4440 0.0226 n 10 6 13 7 36

r,-" (95%)'1" 0.399 0.658 0,306 0.569 0.106

* Values in %:C, with standard error ( ) and calculated correlation coetficient (r~). TT Critical values of r} at 95% confidence, from Arkin and Colton (1963).

are summarized by linear regression in Tables 2 and 3, and selected trends are shown in Figures 1 and 2. Because calculated concentration changes in some major components exceed the total measured concentrations of some minor components, the values are given as the relative percent change of the analyzed value of the daughter sample. That is,

A%~- 100[X,/C,%] (2)

where terms are defined in Eq 1. This normalizes the large differences in weight-percent gains and losses in major and minor components. Mineral densities have been calculated from the mineralogy given by Hower et al. (1976) and Calvert and Klimentidis (1986), but necessarily assumed con-

stant in the other studies. Using ideal mineral densities, the calculations show that rock densities increase slightly with depth, and reveal a systematic error of up to 3% in the volume factor ifa constant density is assumed. This value is small compared to the geologic uncertainties and has no effect on the conclusions to be drawn below.

SiO2.-The small changes, erratic signs, and low r 2 values of ASiO2 with depth or temperature (Tables 2, 3) suggest that silica has low differential mobility. Scatter in individual profiles indicates that changes of less than ± 5% cannot be resolved. The lack of uniform behavior with depth indicates no systematic SlOE changes (± 5%) during diagenesis.

CaO.-Calculated changes in CaO concentration scatter widely, with

v

0J

A % CaO

d - - Weaver & Beck 1971

f _ _ Hower et aL 1976

g ___ Perry & Hower 1970

J . . . . "~ Calvert & k ....... f Klimentidis 1986 n . . . . van Moort 1971

m

S " • / ~"~-

/ m : #

I % i 4 ~ "°

A ~ - s - - [ - - - - - r - - r - - - r - - - r - - 1 , ,

-900 -600 -300 0

A% MgO A %K20

j i . / / ~,

J

B

, , .°° .° -..,~n =='~"T,:

~ " ' . .

%

s D

~ - T - - r - - - r - - . r - -

A %Na20

i

• -60 -30 0 -20 0 20 -80 -40 0 40

FJ<;. l.-Calculated changes in concentrations of selected major oxide components in relative percent (Eq 2) as a function of depth for sample suites in areas with low geothermal gradients and constant TiO}AI203 ratios. Letters identifying profiles are keyed to references in Table I.

DMGENESIS OF SILICICLASTIC ROCKS 353

100 o v

200

300

A % CaO A % Mg20 A % K20 A % Na20

i -

=,r-V-v-r-v=-

H

T'%;;::. I ":':::::

z2_L_

• °°oo

,%. /

\ °%°°°

l . | °°

D

i

t t Muffler & .:, u Z ) White, 1969 v ...... Jennings &

Thompson 1986

w - - Y a u et aL 1988

\ \

N • , , , . N.

° ' " t o \

\ \

/

J

. . r - . . t . . . r - . . r . . . r . . . - - . - - . . . r . -

-200 0 -40 0 40 0 30 60 -40 0 40 80 120 160

FIG. 2.-Calculated changes in concentrations of selected major oxide components in relative percent as a function of temperature for sample suites with high geothermal gradients from the Salton Trough. Letters identifying line types are keyed to references in Table I.

shifts greater than ± 30%, but still show a rather systematic loss of CaO from samples deeper than about 2.5-3 km in Gulf Coast sediments. The data from the Salton Trough show only weak and opposing trends in CaO, but collectively (Table 3, composite) there is a suggestion of CoO loss with depth and temperature.

MgO, FeO(total).-Calculated gains and losses in MgO and FeO(total) show weak and opposing trends. MgO (Fig. 1) and FeO(t) may be lost from Gulf Coast sections, but changes in FeO(0 are opposite in sign and not uniformly strong (Table 2). In the Salton Trough sections, both apparent gains and apparent losses in MgO are present (Fig. 2) and cancel in the composite section (Table 3). Calculated changes in FeO(t) suggest a uniform gain, which is reinforced in the composite section (Table 3).

K20.-The only component consistently added to all mudstones with increasing depth is K20 (Tables 2, 3). Most Gulf Coast sections (Fig. 1) show an increase of up to 20% in K20 content, a trend also present in the three shale sections from Brazil (Fig. 3). All four sections from the Salton Trough also show an increase in K20 of at least 20% in the temperature range 275-300°C (Fig. 2). Although not all sections show statistically strong positive trends with depth (Tables 3, 4), 12 sections show net increases in K,O content of 20--30%. Thus there is little doubt that addition of potassium to mudstones during diagenesis is a general phenomenon.

Nu20.- Calculated changes in Na20 show the greatest scatter and vari- ability. The strongest trends in Gulf Coast mudstones (Fig. 1) suggest a loss of 30--40% Na~O during diagenesis, but not all sections show this trend (Fig. 1; Table 2). In fact, an addition of 10--20% Na20 may exist in the Harris County sections (f, Fig. !), and in the upper 3 km of the Brazilian sections. The overall behavior of Na20 in these sections is not uniform, and no systematic gain or loss through 5 km or 150°C can be identified. An increase of about 30% in Na20 in three of the four sections above 251Y'C from the Salton Trough (Fig. 4) suggests Na20 addition to the higher-temperature pans of these sections.

Volume Changes.- The calculations indicate that the mineral volume of most Gulf Coast sections analyzed decreased by ~ 5% with diagenesis. Although one section (Cameroun) shows an increase, this trend is present

in enough sections to be a general phenomenon in early diagenesis. How- ever, the trends in the Salton Trough are not strong, and volume loss may not be important everywhere.

CORRELATIONS OFF MASS TRANSFER WITH DIAGFENETIC REACTIONS

The progress of many diagenetic reactions is well known to increase with time and depth of burial during diagenesis, whereas temperature defines the kinetics of the reactions (Jennings and Thompson 1986). Many reactions have been identified petrographically (e.g., Burton et al. 1987), but such observations have not been made in most studies listed in Table I. In lieu of petrographic information, changes in modal mineralogy with depth may reveal reactants and products of diagenetic reactions (Ramsay 1973; Hower et al. 1976; Boles and Franks 1979). Changes in mode with depth are available from Calvert and Klimentidis (1986) and can be calculated for five of the Hower et al. (1976) samples by combining the data from their tables 2, 3, and 4. These modes (Figs. 4, 5) show that illite/ smectite, K-feldspar, mica, kaolinite, and calcite all decrease in percentage but increase in size with increasing depth, and that quartz, plagioclase, chlorite, and the percent illite in mixed-layer clays all increase in both wt. percent and size with depth. Illitization of smeclite, dissolution of kaolinite, K-feldspar, and carbonate, and nucleation and growth of chlorite are indicated by these trends (see also Burley and Macquaker 1992); the effects of these reactions on the chemical compositions of the sediments are discussed below.

SM~.L ' I I IE TO ] L U T E ' r ] ~ N S I T I O N : K 2 0 ADDITION

The most universal chemical modification identified is the addition of K20 to mudstones. Because illite/smectite is by far the most important K20-bearing mineral in these rocks, illitization of smectite is the most obvious candidate for this addition. If the potassium precipitated in illite was not all derived locally from the dissolution of delrital feldspars and phyllosilicates, then a correlation should exist between the percentage of illite layers in the clay and the potassium added to the rock. In fact, a

354 ROBERT P. H'INTSCH AND CINDY M. KVALE

5

p ( s ands tone ) ,,,.,,,,,,,,.,,,,,,,,,,

p (shale) ........................

q ( s a n d s t o n e ) . . . . . . . . . . . . i~ , A

q (shale) . . . . . . . . . . . . ~ \ \ • ~, ' \

r ( s ands tone )

r (shale)

-180 ' I ' ' I '

-120 .-60

A % K 2 0

I !

0 60

F,6. 3.-Calculated changes in concentrations of K20 in relative percent for sandstone-shale pairs (p, q, r) as a function of depth from basins offthe northern coast of Brazil (Chang et at. 1986). Shale profiles show systematic gains in K20 in the same depth range where sandstones show losses.

strong positive correlation is observed in all sections where the data are available (Fig. 6), and indicates that illitization of smectite was responsible for addition of K20 to the mudstones (as predicted by Aagaard et al. 1990 and found by Awwiller 1993). For most sections, illitization of smectite extrapolated to 100% illite would involve an increase of ~ 35% K20 relative to the reference sample. However, the reference samples (at about 20% illite) may already have gained some K20, and the original sediments may have gained as much as 50% K20 (Jennings and Thompson 1986) if the parent clay was > 90% smectite. These results show that dissolution

of detfital K-feldspar and mica alone cannot provide the K20 required by illitization of smectite, as proposed by Hower et al. (1976).

There is debate over the mechanism of illitization of smectite, either by ion exchange (e.g., Hower et al. 1976; Boles and Franks 1979), with the implication that the mechanism of illitization is a pseudomorphic one, or by neoformation (e.g., Nadeau and Bain 1986; Burley and Flisch 1989; Freed and Peacor 1992; Small et al. 1992a, 1992b). The modal and grain- size information in the Harris County section (Fig. 7) shows that the smaller size fraction decreases in wt. percent, and that the illite/smectite in the larger size fractions actually increases (see also Glasmann et al. 1989). This coarsening suggests that the reaction mechanism involves (1) dissolution of the finest fraction of the illite/smectite (as also suggested by Ahn and Peacor 1986), and (2) neoformation or overgrowth of new illite/ smectite on the larger (coarsening) grains. Overgrowths further open the possibility that other dissolving mineral grains (kaolinite, feldspars) may have provided some of the required aluminum for the new illite layers. Some degree of Ostwald ripening may also be involved (Eberl and Srodon 1988).

The very strong indication that K20 is added to smectite-bearing rocks during illitization requires the introduction of K20 via the diagenetic fluid. This inference is supported by the very common depletion of K + from diagenetic pore fluids (Collins 1969; Manheim and Bischoff 1969; Sayles 1979; Hanor 1987; Land et al. 1988; Mottl and Gieskes 1990). The source of the potassium in the fluid is not known, but the calculated decrease in K,O in sandstones parallel to the calculated increase in K20 in mudstones in all three Brazilian basins (Fig. 3) suggests that dissolution of potassium- beating minerals (alkali feldspars) in the sandstones (see also Boles and Franks 1979; Awwiller 1993) provides the potassium that is ultimately fixed in the shales. Although all three basins show the same trend, the data are sparse, and a much more detailed analysis ofsandstone-mudstone pairs is needed to test this hypothesis.

SiO 2 MOBILITY

Contrasting opinions on the differential mobility of silica exist in the literature. Boles and Franks (1979) proposed that the SiO2 released from smectite during illitization is transported into more permeable sandstones, where it precipitates as overgrowths. They predicted a Joss of 4-5% SiO2 in mudstones across an interval several km thick and over a temperature range of ~ 60-20(YC. Alternatively, released silica may precipitate locally in new silicates. Calvert and Klimentidis (1986) cited fine-grained, subhedral quartz in the deeper samples of the Kenedy County sections as evidence that quartz cements in both sandstones and shales were self- sourced.

Two tests of these hypotheses can be made. First, the calculations here do not show any systematic change in SiP2 (Table 2), suggesting that silica

TABLE 4.-- Calculated gains and losses (wt %) of components in smectite/illite and volume factors"

Depth (kin) Samples S~; TiP, AI_,O~ FeO MgO CaP Na:O K20 Fv (hi)

1.85 a 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 1.00 2.15 b 2.98 - 0 . 0 2 0.00 0.17 0.29 1.13 0.16 0.47 1.06 2.45 c 1.17 0.02 0.00 - 0.31 0.82 0.64 - 0.08 0.84 1.04 2.75 d - 2.23 0.07 0.00 - 1.18 0.46 0.77 - 0. l 7 0.97 0.98 3.10 e - 2 . 9 7 - 0 . 0 4 0.00 - 0 . 4 3 0.10 - 0 . 0 9 - 0 . 3 0 0.71 0.97 3.40 f - 6 . 4 2 - 0 . 0 4 0.00 - 1.21 - 0 . 3 0 - 0 . 7 2 - 0 . 6 2 1.12 0.92 4.00 h - 9.50 - 0 . 0 4 0.00 - 2.65 - 1.02 - 0 . 9 9 - 0 . 5 7 3.01 0.88 4.30 i - 10.32 - 0 . 0 4 0.00 - 3.19 - 1.21 - 1.04 - 0.66 2.93 0.87 4.60 j - 14.09 - 0 . 0 4 0.00 - 3 . 2 5 - 1 . 1 5 - 0 . 8 6 - 0 . 7 7 2.19 0.81 4.90 k - 12.81 - 0 . 0 5 0.00 - 3.30 - 0 . 9 8 - 0 . 8 7 - 0 . 7 3 2.65 0.84 5.20 I - 11.46 - 0 . 0 4 0.00 - 2.55 - 0 . 8 4 - 1.07 - 0 . 7 0 2.61 0.86 5.50 m - 13.45 - 0 . 0 6 0.00 - 3.40 - 1.14 - 1.01 - 0.67 2.41 0.84

* Calculated from analyses given in Hower et al. 1976, assuming A1203 immobility and a starting composition of sample a; sample g (3.7 km) not analyzed. t Sample depth label for reference in Figures 6 and 8.

wt. % of total

v

K-feldspar (2->10)

0

I. 20 40 60 80 100

I I I I I I I I I I

Illite-Smect

!

DIAGENESIS OF SILICICLASTIC ROCKS 355

0.1-0.5

Fie. 4.-The variation in mineralogy (wt %) as a function of depth in the Harris County, Texas, section, calculated from Hower et al. (1976). Selected fields are subdivided to indicate the wt % contribution of the different size fractions. Note that the percentage of minerals in the > 2 ~m sizes increases for most minerals, even if the total percentage drops with increasing depth.

is not lost uniformly from smectite-bearing mudstones. This is surprising in that illitization of smectite does occur in all the wells sampled, and its effect on K:O is pronounced (see above). Although the ~ 5% loss of SiO2 may be undetectable at the resolution of the calculations, these results clearly do not support systematic and significant upward transport of silica to shallower depths.

A second test lies in modal quaaz in mudstones. If modal quartz is constant with illitization, then silica transport is implicated. If modal quartz increases, then released silica probably precipitated locally. In fact, modal quartz does increase significantly with depth in the Hams County section, especially in the > 2 tzm fractions, as well as in the Kenedy County sections (Calvert and Klimentidis 1986), suggesting local silica precipitation. The fact that the increase is mainly in the 2-10 vm fraction (Fig. 4) further implies that grains < 2 tzm do not serve as nuclei for overgrowth, which may relate to the relatively high surface energy/mole of these smaller grains, following Ostwald ripening (Eberl and Srodon 1988).

The local precipitation of quartz suggested by these increases is consistent with the ability of quartz to buffer the aqueous activity of silica at temperatures as low as 50"C and in most cases by 100*C quartz, as dem- onstrated using stable isotopes (Land 1984; Yeh and Savin 1977) and aqueous silica concentrations (Kharaka and Berry 1974; Kharaka el al. 1977; Merino 1975; Arnorsson 1975). Thus the SiO2 released from illi- tizalion should be precipitated locally as quartz at depths as low as 2-3

km. From our calculations, and from the above arguments, SiO2 transport out of mudstones is probably not significant at the scale of kilometers.

ALBITIZATION OF PLAGIOCLASE

Intermediate plagioclase (An2o-An3o) is ubiquitous in the Gulf Coast sediments (Land 1984) and probably was present in all sediments considered in this study, but because An content cannot be identified easily by XRD methods, none of the studies include the composition of the plagioclase present in the rocks. Although intermediate plagioclase in sandstones is generally albitized (Browne and Ellis 1970; Land and Milliken 1981; Boles 1982; Land 1984; Walker 1984; Fisher and Land 1986; Gold 1987; Saigal et al. 1988; Aagaard et al. 1990; Burley and Macquaker 1992; Ramseyer et al. 1992), the only report ofalbitization in the sections studied is from Calvert and Klimentidis (1986), who noted subhedral authigenic albite from Kenedy County mudstones.

Our calculations ffables 1, 2; Figs 1, 2) show variable changes rather than pervasive increases in Na20, suggesting that albitization does not lead to fixation of Na20 from diagenetic fluids. This is consistent with the compositions of subsurface pore fluids that show no important changes or only small decreases in the Na+/CI ratios relative to sea water (Hanor 1987). Addition of NazO to the Harris County section (Fig. 4) occurs in spite of the systematic loss of Na20 from the reacting smectite in these rocks (Table 4). Supporting evidence for albitization in the Harris County

356 ROBERT P. ItTXISCH AND CINDY M. KI:)ILE

samples is the increase in size as well as in abundance of plagioclase with depth (Fig. 4). Here albitization occurs between about 60--90°C, at the low end of the temperature range reported by Aagaard et al. (1990), and lower than the 100-150°C reported by Boles (1982) and Fisher and Land (1986) in adjacent areas. In summary, it appears that albitization can increase the concentration of Na,O in some mudstones, but pervasive increases are not revealed by these calculations.

CHLORITE PRECIPITATION

Several studies have documented the growth of chlorite during diagenesis (e.g., Curtis et al. 1984; Chang et al. 1986; Burley and iacquaker 1992). In particular, in Harris County Hower et al. (1976) document an increase in the wt. percent and grain size of chlorite with depth (Fig. 4), which is strong evidence for chlorite precipitation during diagenesis. At least some of this chlorite appears as thick packets intergrown with smectite (Ahn and Peacor 1985), suggesting that this chlorite nucleated and grew during illitization of smectite. Chang et al. (1986) have identified an anal- ogous process of chloritization of saponite. Precipitation of chlorite directly from solution has also been observed (Burton et al. 1987). However, the increase in grain size of chlorite with depth in the Harris County samples also suggests an overgrowth mechanism.

Precipitation of chlorite from solution might be expected to lead to addition of FeO and MgO to the rock; this has been proposed to account for the depletion of magnesium in formation waters of western Canada (Hitchon et al. 1971). Our calculations, however, show no systematic increases in either MgO or FeO(t) with depth. Only the Cameroun section shows an increase in MgO (Table 2), but Dunoyer de Segonzac (1964) interpreted this chemical trend as one inherited from deposition rather than resulting from diagenetic reactions. The possible covariance of these components with CaO (see below) suggests that small changes of FeO(t) and MgO could be masked by the variable carbonate contents of the rocks. Thus with the data available the growth of chlorite cannot be shown to lead to an increase in the concentrations of FeO(total) and MgO; the apparent source of these components is dissolving smectite, carbonate, and probably iron oxides.

CALCITE DISSOLUTION

CaO.-Hower et al. (1976) and Calvert and Klimentidis (1986) document loss of calcite and CaP with depth. Most Salton Trough sections also show this loss. Although illitization of smectite (Table 4) and dissolution of plagioclase may also contribute to CaP loss, the losses in the Gulf Coast sections are totally overwhelmed by calcite dissolution, which apparently is a regional phenomenon.

MgO, FeO (toNI).-Most sedimentary carbonates in mudstones contain up to 5 mole % MgO and FeO (Muffler and White 1969; Boles 1978; Land 1984; Yau et al. 1988). If dolomite or ankerite do not precipitate, then calcite dissolution might be expected to produce losses of MgO and FeO as well as of CaP. In some sections there is a moderate to strong covariance of changes in CaP and MgO concentration (especially in the Galveston section, Fig. 7), suggesting that their changes are related to carbonate dissolution. Supporting evidence for this comes from the temperature dependence of the Ca/Mg ratio of pore fluids, implying a composition buffered by Ca- and Mg-bearing minerals (Land et al. 1988). If dissolution of calcite locally increases the activities of Ca + + and Mg + + in the adjacent fluid, then some ion exchange with the smectites would be expected. In- deed, the CaP and MgO contents of the shallower smectite samples of Hower et al. (1976) (samples B-F) do increase, suggesting at least local influence of carbonate dissolution on water composition.

Boles (1978) and Boles and Franks (1979) argue that the CaP, MgO, and FeO released in illitization of smeetite are transferred to adjacent sandstones, where they precipitate as carbonate cements. The overall loss

3

v

wt % of total

3

K

o 20 40 60 80 100

Ft~. L-Variations in mineralogy (wt %) with depth in the Kenedy County sections, calculated with X-ray modal data from Calvert and Klirnentidis (1986). A) a composite profile produced by averaging modes of samples from narrow depth intervals (J, Table 1). B) modal changes in a single well (K, Table 1).

of CaP from most mudstones is consistent with this hypothesis. However, the above calculations do not show any systematic gains or losses of FeO(total) or MgO, with the implication that any iron or magnesium released by calcite dissolution is taken up by concurrent reactions, probably chlorite precipitation.

EFFECT OF DIAGENETIC REACTIONS ON BULK ROCK COMPOSITIONS

It is well known that diagenetic reactions strongly modify the mineralogical and chemical compositions of some rocks, and that even just selective dissolution of feldspar grains may change inferences of provenance made from the proportions of quartz, feldspar, and lithic fragments (e.g., Milliken 1988; Milliken et al. 1989). In this section we show the effect of potassium and calcium metasomatism identified here on the bulk

DL4GENESIS OF SILICICLASTIC ROCKS" 357

60

50

40

30-

~ 10

0

-10

-20

-30

f A,nM Hower et al 1976 |

J 0 Calvert & Klimenfidis 1986 •

k • Calvert & Klimentidis 1988

[9 • Chang etal. 1986 q ~ Changefal. 1986 |

V • jennings & Thompson 1986 A k _ _ _ _ _ " 1 •

© F • c • ltz It S t I

r ? T 'I r

20 40 60 80 100

% Illite in I /S

FIG. 6.--Variation in the calculated change in K20 content in a mudstone with percent illite in the illite/smectite mixed-layer clay as reported in the original sources (Table 1). The positive correlation in all wells is strong evidence that conversion of smectite to illite is responsible for the increase in K20 in the host rock.

compositions of mudstones by superimposing the compositions of mudstones on sandstone classification diagrams (Fig. 8).

Na20/KzO.- One of the chemical variables commonly used in classification of sedimentary rocks is the Na20/K20 ratio of the rock (Englund and Jorgensen 1973; Schwab 1978; Roser and Korsch 1986; Argast and Donnelly 1987; Pettijohn et al. 1987, p. 57). Illitization and albitization could have important but opposite effects on this ratio, as shown by the data from Harris County in Fig. 8A. Addition of Na20 to the rock in the five shallowest samples (A-E, in Figure 8A) moves the sediment composition from the lithic arenite field into the graywacke field. In deeper samples, illitization of smectite moves the composition back across the lithic arenite field into the arkose field. From E to M diagenetic reactions decreased the Na/(Na + K) ratio by 30%. Thus the Na/(Na + K) ratio of a rock may be as strongly affected by diagenetic reactions as by provenance, and should be interpreted cautiously.

Cad.-Dissolution of carbonate, as expected, also has a large effect on rock composition. Again in the Harris County example, Figure 8B shows

30

O

<1

0

-30 -'

-60 "

-90 , i i

-800 -600

a l l

•l lt l g.

e , Q ,

-400 -200

& wt. % Ca P

F~. Z-The relationship between A% MgO and A% CaP in the Galveston sections (Perry and Hower 1970). The positive correlation suggests that MgO may be removed from some sections as magnesium-rich calcite is dissolved.

that dissolution of calcite moves bulk compositions from the calcareous graywacke field to the arkose field, with a change of 40% in the (Na + Ca)/(Na + Ca + K) ratio. Subtraction of the calculated K20 added during illitization moves the rocks in the arkose field to the graywacke field, leaving the change in composition from calcite dissolution still to 35 mole %. Clearly, dissolution of carbonate may have a major impact on the bulk composition of some sedimentary rocks, turning marly rocks into mudstones, and with a possible increase in porosity.

Reservoir Sandstones.-In spite of the important effect of diagenetic reactions on mudstone composition, they may be small relative to the chemical changes occurring in some sandstone reservoir rocks. Very significant albitization, iUitization, feldspar dissolution, and carbonate, quartz, and laumontite cementation are well documented in some reservoir sandstones (Boles 1978; 1982; Boles and Flanks 1979; Merino 1975; Land el al. 1987). These mineral reactions would likely have a larger effect on sandstone composition than that documented here in mudstones and clearly require at least local transport of all components, including alumina. The contrast between the small chemical changes calculated here and the larger chemical changes observed in sandstone reservoir rocks during diagenesis is probably related to the dominance of organic acids in reservoir sands rather than chloride (Fein 1991), but the differences may also be due to different scales of transport, to rock type, or to different scales of sampling. Our results show no important transport in mudstones of many components on the scale of about 0.5 km in sections in which reservoir sandstones are present. We must conclude that the transport indicated by the study of reservoir sandstones occurs on a scale smaller than 0.5 km, and has no detectable effect on kilometer-scale vertical trends of mudstone composition.

APPLICATION TO ANCIENT ROCKS

These observations bear directly on comparisons of modern and ancient sedimentary rocks as well as on comparisons of rocks from different tectonic settings (e.g., Schwab 1975, 1978). Because CaP and K20 contents of sedimentary rocks are so sensitive to the progress ofdiagenetic reactions, the compositions of sediments and sedimentary rocks should be compared only if the progress of diagenetic reactions is similar in all rocks. Indeed, the increase in K20/Na20 ratio in pre-Mesozoic sediments relative to Mesozoic + Cenozoic sediments described by Pettijohn et al. (1987, p. 507) could be an example of this. Diagenetic reactions tend to be more

358 ROBERT P. 14 7NTSCH AND CINDY M. KVALE

1.0

0 . 9 - -

g.

• 0.8 -

< + 0.7 - .~=.~

0 . 6 -

0.5

1.0

0.9

0.8 O

+ 0 . 7 ,p,,i

SUBLITH1C ARENITE ~ quar tz / . . , . ~ . ' , .

0 . 6 - -

0.5

\

4

f e d O ~ 4 1 - - ' - - ' - ' - - " - - " ' ~ .x . . smecti te

• i h ~ / ~ c a TITE - ~ m 7 iLLiTiZATION OF S MEC , Mudstones ~1 x O i l l i te /smect i te

xx x j i l l i t e X, x i l l i te /smect i te

muscovi te X

0 0.2

kaolinite ,/ ~ paragoni te " ' ,d I I I I I I

0.4 0.6 0.8 1.0 N a / N a + K ( a t o m %)

quar tz /

. . . . .. • : ' i ' / ' : ' . : :.'..'.: - " : : ',":'..' :..'2;.;~:,~:=2;;,:.'.".: . " : : ":i':'.: :.'."

v d -

~ , ~ k --TZZATION OF sMEc~TE J ILLI

, ~ muscovi te / x N I I I I I I

0 0.2 0.4 0.6

kaolinite

N a + C a / N a + C a + K (atom

I Mudstones

O i l l i te /smect i te

paragoni te ~ai I I I

0.8 1.0 % )

FIc. 8.-Compositions of mudmcks (filled circles) and smectites (open circles) from Harris County, Texas, reported by Hower et al. (1976) superimposed on sandstone chemical classification diagrams. A) plotted in terms of Si/Si + AI and Na/Na + K (atomic proportions); adapted from Pettijohn el al. (1987, p. 553). B) plotted in terms of Na + Ca/Na + Ca + K and Si/Si + AI; adapted from Garrels and McKenzie (1971, p. 213). Letter symbols are keyed to the rock depths in Table 4 (where G is a sample from 3.70 km depth). Arrows indicate chemical trends with increasing depth, or progress of diagenesis. The increase in Na in the shallower rock samples A-E (Fig. 8a) is in spite of an increase in the K content of illite/ smectite samples a-e, and is probably caused by albitization of plagioclase. Here albitization appears to move the rocks from the litharenite to the compositional graywacke field. The overall decrease in the Na/Na + K ratio of the rocks is probably caused by the dominance of illitization of smectite over albitization, which moved the rocks from the compositional graywacke field into the arkose field. The calculated compositions of end-member illite and smec- rite (X) based on the results of Hower et at. (I 976) and the analyzed compositions of illi"te (x) from Boles and Franks (1979) are included for comparison. The dramatic effect of calcite dissolution on rock composition moves the rocks from the compositional graywacke field to the compositional arkose field. For further explanation, see text.

complete in older rocks, and comparisons of the alkali contents of rocks of differing ages and progress of reaction should be made with caution.

M E C H A N I S M S D ~ V l N G R E A C T I O N S

The large changes in K20 and CaO calculated above require large fluxes of diagenetic fluids through the mudstones. Lack of apparent change in the other components is consistent with buffering of the fluid composition by the mineral assemblage in the rock, as shown in many rocks (Yeh and Savin 1977; Land 1984; Kharaka and Berry 1974; Kharaka et al. 1977; Merino 1975; Arnorsson 1975), such that rates of dissolution and precipitation are equal and neither fluid nor rock changes composition signifi-

cantly. If infiltration of a fluid far from equilibrium with the rock can be eliminated as a driving force for reaction, then this force must have come from within the rock itself. Low-temperature solid-solution relationships in plagioclase (Smith and Brown 1988) show that intermediate plagioclase is metastable with respect to albite and other calcic minerals (see also Ramseyer et al. 1992); this thermodynamic metastability of intermediate plagioclase enhances its solubility and helps drive albitization. The universal illitization of smectite is indirect evidence that metastability of smectite also drives illitization (especially with increasing temperature), but the thermodynamics of illite/smectite solid solutions are not known in enough detail to prove this.

In the context of mineral assemblages buffering fluid compositions, the

DIAGESESIS OF SILICICL4STIC ROCKS 359

I I I I I __ l I I - - ¢ I --

- - ' - - ' ~ - ' - - ' - - - ' CM-a++~--g ++

- £ I - - K + + S m e c t i t e ~ l l l i t e + , H 4 __ _ . . . - -- ~_'

- - -- .-~: ~ " - - ~ - ~ H % C a l c i t e - - I I ~ I - - I - - [ I - - ~ ' - - I - - =

I ± , t S ( d i ) , - s s o l u t i o n _ _ ,

- 1 - - - ' ~ - ' . - . ' ~ . . ' ~ : 5 / : . . . . . ' .

~ - ' - - - - -= :., - - - - - - ~ , 7:. -: T-== :.- . :. . :i P.: ; ; d o l o m i t e : ;..;.i..: :.:...:. ::": i..,":

]-:...: ( ,: . ~... : : , : ] : .~, :....) ( r e l e a s e d ) i : : : , . " . : : : / : . ~ . . " ~ '..~.".: : / :....: : . , : : : ( i : : ::i:

i :"."4.."~::. " : , . " " :~ ..". : . :: ! . : : ) . . " ..". ::. : : . - " . . - . i :. : . . " - . I - . " . : : . : : : . : - - " " - ± :.: H + K-feldspar:: :~::.".:.....:..::.:.::/."v.~-~-~-::--~:: :; • ""' ~ m s s u m u u j ' ~ : - - - ' - ' : - n " ' ' ' : ' : ' ~ ' - -

. . . . • , . . . . , - - , - , a + + _ ~ M g + + ~ ~ : . I I - I

K ÷ + S m e c t i t e : I l l i t e + H ÷=- ~ l , - - - l , ~ , - - - i ~ (ppo .~_ , - -- H + + C a l c i t e ~- - -

l l - - ~",,..._ _..,--¢' I

! -ZL f - ( d i s s o l u t i o n ) L ' _

- , _ , - - K + ~ ,~Mg++~Ca++_~--~H20~---~ I~ i -- i ''" i

Fro. 9.-Schematic diagram summarizing diagenetic changes in siliciclastic rocks that change the composition of the rocks. The thermodynamic metastability of smectite drives illitization, which fixes K20 in the mudstone (dashed and wavy lines) and releases H + . This acid dissolves carbonate in calcareous layers (&), which lowers the concentration of CaO in the mudston¢. It is very likely that the fluids react with any sandstones into which they may rise (stippled region, center), where they may lose CaO through carbonate precipitation and gain K20 through K-feldspar dissolution. Upon further rising, they reenter mudstones, where illilization is again important. See text for other details.

dissolution of carbonate in essentially all Gulf Coast sections (see also Lundegard and Land 1986) is surprising. Because calcite is ubiquitous in the stratigraphic sections studied, diagenetic fluids should already be sat- urated with respect to calcite, and there should be no net dissolution or precipitation of calcite. Nevertheless, the nearly universal dissolution of calcite implies an increase in pH (Garrels and Christ 1965, p. 88) and destruction of solid-fluid equilibrium. However, precipitation of K20 in diagenetic illite decreases the pH, an example of a "reverse weathering" reaction of Mackenzie and Garrels (1966); to maintain equilibrium, the remaining assemblage would react with the brine to reestablish a buffered composition. In the absence of detailed information on reaction mechanisms, mineral compositions, and aqueous and organic complexes, with- out which pH cannot be calculated, only this qualitative relationship between these two reactions can be proposed:

smectite~, + 2KCI = illit% + 2HCL (3)

+ 2HCI + calcite = CaCI2 + H2CO3 (4)

2KCI + smectite, + calcite = illit% + CaCI2 + H2CO~ (5)

where the subscript ss refers to solid solution. In a closed system, the progress of Reactions 3 and 4 is limited by the increase in concentration of the aqueous products. However, in an open system, Reaction 5 can preserve local solid-fluid equilibrium, and can be limited only by introduction of K + and removal of Ca + +. Reaction 5 can be driven by the metastability of smectite with increasing temperature that leads to illitization (Reaction 3). lllitization decreases the local pH, which is rebuffered by the dissolution of carbonate. Equally, dissolution of a small volume of calcite with increasing temperature increases the pH, and because the percentage of illite in illite/smectite varies directly with pH (Helgeson and Mackenzie 1970; Small et al. 1992a), illitization and the fixation of K20 is the consequence. Because mudstones are shown here to act as sinks for K + as well as sources of Ca + + (see also Boles and Franks 1979), illitization may be the in situ "acid-producing" reaction that could allow dissolution of nannofossil calcite to buffer local pore waters and account for the modal and isotopic data of Lundegard el al. (1984), Lundegard and Land (1986),

and Land et al. (1987) that su~ests that the carbonate in cements in sandstones was derived from underlying mudstones.

A WORKING MODEL OF DIAGENESIS IN S ~ C l ' l r l ' ~ - ~ R ~

The above calculations and inferred reactions can be drawn together in the scenario of Figure 9. In actively compacting sedimentary sections, diagenetic aqueous fluids are slowly expelled upward through the sedimentary section on a scale greater than several kilometers (Land et al. 1987; Wintsch and Kvale 1991). Detrital smectites in mudstones react with K + from this diagenetic fluid to precipitate illitic mica (Reaction 3), and increase the KzO content of the mudstone (Tables 2, 3). This reaction produces H ÷ that reacts with detrital carbonates in the mudstones to dissolve them (Reaction 4), which causes a loss in the CaO content and possibly the MgO content of the mudstone (Tables 3, 4). The net reaction (5) occurs in the mudstone, but it may or may not occur in every hand- specimen-size volume of rock, depending on the mineralogy of that rock. As the fluid rises, it may leave the mudstone and enter a sandstone (Fig. 9). If that sandstone contains more CO2 than the mudstone, the CaC12 carried in that fluid may react with dissolved carbonate to precipitate calcite or dolomite (Reaction 4 in reverse). This releases H +, which then reacts with any metaslable detrital feldspars that are in the sandstone• If K-feldspar is present, then the sandstones lose K~O (Fig. 3) to the diagenetic fluid, and if the feldspar is albitized, removal of the K + is required (Aagaard et al. 1990). The fluid that rises into overlying mudstones is then relatively depleted in Ca + ÷ and enriched in K +, and the process repeats (Fig. 9).

This scenario incorporates all observations of changes in rock and fluid composition identified and summarized above, as well as mineralogical, modal, and textural observations, and data from the literature. Clearly what is needed is a single, integrative study incorporating all of these kinds of data in a single stratigraphic section. With such data many details of solid-fluid reactions and reaction mechanisms could be identified, and quantitative mass-balance calculations including fluxes of components "flushed" from sedimentary basins could be attempted. At this point the model of Figure 9 awaits further testing.

D1AGENESIS OF SILIC1CLASTIC ROCKS 361

HIICflDN, B., BUUNGS, G.K., AND Ktov~r~, J.E., 1971, Geochemistry and origin of formation waters in the western Canada sedimentary basin-IlL Factors controlling chemical composition: Geochimica el Cusmochimica Acta, v. 35, p. 567-598.

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Received 17 June 1991; accepted 6 October 1993.

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Differential Mobility of Elements in Burial Diagenesis of Siliciclastic Rocks